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CARNEGIE
INSTITUTION

Annual Report of the Director
Geophysical Laboratory

5251 BROAD BRANCH ROAD, NORTHWEST, WASHINGTON, D.C. 20015-1305

1990-1991

For the year July 1, 1990-June 30, 1991

Issued December 1991

Papers from the Geophysical Laboratory

Carnegie Institution of Washington

NO. 2250

Digitized by the Internet Archive

in 2012 with funding from

LYRASIS Members and Sloan Foundation

http://www.archive.org/details/annualreportofd199091carn

Geophysical Laboratory

Washington, District of Columbia

Charles T. Prewitt
Director

Published by: Geophysical Laboratory

5251 Broad Branch Rd., N.W.
Washington, D.C., 20015-1305
USA

ISSN 0576-792X
December 1991

When used in bibliographic citations, The Annual Report should be cited as follows:

Author, Title, Annu. Rep. Director Geophys. Lab., Carnegie Instn. Washington, 1990-1991, pagina-
tion, 1991.

GEOPHYSICAL LABORATORY

Contents

Introduction 1

Igneous and Metamorphic Petrology — -
Field Studies 3

Global Convection and Hawaiian Upper Mantle

Structures. T.Neil Irvine 3

Megacrystalline Dunites and Peridotites: Hosts

for Siberian Diamonds. N. P. Pokhilenko, D.

G. Pearson , F. R. Boyd,

andN.V.Sobolev 11

Mantle Metasomatism: Evidence from a MARID

- Harzburgite Compound Xenolith.

F.R.Boyd 18

Boron and Beryllium Concentrations in Subduc-

tion-Related Metamorphic Rocks of the

Catalina Schist: Implications for Subduction-

Zone Recycling. Gray E. Bebout, Jeffrey G.

Ryan, and William P. Leeman 23

Laser Fluorination of Sulfide Minerals with F2

Gas. D. Rumble, J. M. Palin,

andT. C. Hoering 30

Stable Isotope and Trace Element Indicators of

Devolatilization History in Metashales and

Metasandstones. Gray E. Bebout 34

The Fa Content of Normative ol.

Felix Chayes 40

Igneous and Metamorphic Petrology —
Experimental Studies 45

Raman Spectra of High-Temperature Silicate
Melts: Na20-Si02, K20-SiC>2, andLi20-SiC>2
Binary Compositions. John D. Frantz and
Bjorn O. Mysen 45

Peralkalinity and H2O Solubility Mechanisms in
Silicate melts, Bjorn Mysen 53

Partitioning of fluorine and chlorine between apa-
tite and non-silicate fluids at high pressure and
temperature. James Brenan 61

Investigation of Fluid Immiscibility in the System
H20-NaCl-CC>2 Using Mass Spectrometry and
Microthermometry Techniques Applied to
Synthetic Fluid Inclusions.
Robert K. Popp,John D. Frantz,
andThomas C. Hoering 68

Akermanite-Gehlenite-Sodium Melilite Join at

950 C and 5 kbar in the Presence of CO2 +

H2O. H.G. Huckenholz, H.S. Yoder, Jr., T.

Kunzmann, and W. Seiberl 75

Merwinite Stability and High-Temperature Phase

Relations in the Presence of CO2 +H2O. H. G.

Huckenholz, H. S. Yoder, Jr.,

and W. Seiberl 81

The System Mg2Si04-Fe2SiC>4 at Low Pressure.

Hiroko Nagahara, Ikuo Kushiro,

andBjorn O. Mysen 88

Fe3+, Mg order-disorder in heated MgFe204: a

powder xrd and 57pe mossbauer study. H. St.

C. O'Neill, H. Annersten andD. Virgo ....93

Crystallography - Mineral Physics 101

Predicted High-Pressure Mineral Structures with
Octahedral Silicon. Robert M. Hazen and Larry
W. Finger 101

Simultanous High P-T Diffraction Measurements
of (Fe,Mg)Si03-Perovskite and (Fe,Mg)0
Magnesiowiistite: Implications for Lower
Mantle Composition. YingweiFei, Ho-Kwang
Mao, Russell J. Hemley, and Jinfu Shu ..107

High-Pressure Crystal Chemistry of Iron-Free
Wadsleyite, p-Mg2SiC>4 Jinmin Zhang, Rob-
ert M. Hazen, and Jaidong Ko 115

Phase Transitions in Framework Minerals.
David Palmer 120

First-prirciples Studies of Elasticity and Post-
Stishovite Phase Transitions in Si02.
Ronald E. Cohen 126

Molecular Dynamics Simulations of Melting of
MgO at High Pressures. Zhaoxin Gong,
Ronald E. Cohen, and Larry L. Boyer ... 129

Glass Diffraction Measurements with Polychro-
matic Synchrotron Radiation. Charles Meade
and Russell J . Hemley 135

X-Ray Diffraction of Solid Nitrogen-Helium Mix-
tures. Willem L. Vos, Larry W. Finger,
Russell J. Hemley, Ho-Kwang Mao,
Jing Zhu Hu, Jin Fu Shu, Richard LeSar,
Andre de Kuijper,
and Jan A. Schouten 138

CARNEGIE INSTITUTION

Evidence for Orientational Ordering of Solid Deu-
terium at High Pressures. Russell J. Hemley
and Ho-Kwang Mao 141

BlOGEOCHEMlSTRY 147

Nitrogen Isotope Tracers of Atmospheric Deposi-
tion in Coastal Shelf Waters off North Caro-
lina.
Marilyn L. Fogel and Hans W. Paerl 147

Nitrogen Diagenesis in Anoxic Marine Sediments:
Isotope Effects. David J. Velinsky, David J.
Burdige, and Marilyn L. Fogel 154

The Isotopic Ecology of Plants and Animals in
Amboseli National Park, Kenya. PaulL. Koch,
Anna K. Behrensmeyer,
and Marilyn L. Fogel 163

Rapid Racemization of Aspartic Acid in Mollusk
and Ostrich Eggshells: A New Method for
Dating on a Decadal Time Scale. Glenn A.
Goodfriend, David W. von Endt,
and P.E. Hare 172

A Burning Question: Differences between Labo-
ratory-Induced and Natural Diagenesis in Os-
trich Eggshell Proteins. A S. Brooks, P.E. Hare,
J.E. Kokis and K. Durana 176

Publications 181

Personnel 185

GEOPHYSICAL LABORATORY

Introduction

Last year's introduction to the Annual
Report described the co-location of the
Geophysical Laboratory and the Depart-
ment of Terrestrial Magnetism in our new
and newly-renovated building complex on
Broad Branch Road. This year the con-
struction and renovation were completed
and both departments are pursuing their
normal research objectives with almost all
of our equipment operating as it should.
Moving and the relocation of laboratories
has been disruptive for many staff mem-
bers, but we all believe that the new envi-
ronment and proximity to colleagues at
DTM are worth the effort.

The principal new research initiative
for the Geophysical Laboratory this year is
our participation in the new Center for
High Pressure Research with the State
University of New York at Stony Brook
and Princeton University. This is one of the
14 Centers established in 1991 through
funding by the NSF Science and Technol-
ogy Center Program. Depending on Con-
gressional budgetary approvals, NSF in-
tends to continue the Program for at least
four more years, and it is possible that
funding could extend for a total of eleven
years. The funds supplied by NSF together
with contributions from the three institu-
tions and additional external grants will be
used to support a variety of initiatives re-
lated to high-pressure research.

The Center is composed of staff, students,
and laboratories at the Geophysical Labo-
ratory, the Department of Earth and Space
Sciences at Stony Brook, and the Depart-

ment of Geological and Geophysical Sci-
ences at Princeton. GL staff members
involved with the Center are Francis Boyd,
Ronald Cohen. Larry Finger, Robert Hazen,
Ho-kwang Mao, Russell Hemley, Bjom
Mysen, Charles Prewitt, and David Virgo.
In addition to the above institutions, col-
laboration is being developed between the
Center and other laboratories in universi-
ties, industry, and government. The princi-
pal goal of the Center is to study fundamen-
tal questions about the Earth's evolution,
structure, and dynamic state, and about the
nature of interiors of other planets. In addi-
tion, we expect to generate extensive new
information about material properties at
high pressures and temperatures, and to
synthesize new materials of interest to phys-
ics, chemistry, and materials science as
well as to the earth sciences. Experimental
work will be complemented by theoretical
computer simulations and by development
of new equipment and techniques for high-
pressure research, including larger-volume
experimental apparatus.

Systematic high-pressure work on ma-
terials of geological interest has been a
fundamental component of Geophysical
Laboratory activities since 1904 and GL
staff have played the major role in the
development and utilization of many types
of high-pressure apparatus, including pis-
ton-cylinder, cold-seal, gas-media, and dia-
mond-anvil devices. For example, in re-
cent years Ho-kwang Mao and his col-
leagues have led the development and ap-
plication of the diamond-anvil cell for ex-

CARNEGIE INSTITUTION

perimental studies at high pressure. Static
pressures of about 5.0 megabars — substan-
tially greater than that at the Earth's cen-
ter— have been attained, the important pres-
sure-measuring scale using ruby fluores-
cence has been extended to 1.8 megabars,
and pressure-scale x-ray diffraction studies
have been extended above 3 megabars.
Techniques have been developed for heat-
ing samples within the cell by laser and for
studying them by means of a variety of
spectroscopic techniques. The establish-
ment of the Center will allow us to continue
and extend this kind of innovation to greater
extremes of pressure and temperature, larger
sample volumes, and experiments on many
different kinds of materials.

In recent years, Mao, Russell Hemley,
and colleagues have have concentrated
much of their work on inert gases that
crystallize into solid forms at high pres-
sures. Their experiments with solid hydro-
gen have exceeded 2.5 megabars, where
they discovered new phase transitions and
the first evidence of transformations into
the metallic form. To support this research,
a number of staff members and postdoctoral
fellows have been active in developing and
using x-ray and infrared beam lines at the
National Synchrotron Light Source,
Brookhaven National Laboratory, for ex-
periments that could not be performed sat-
isfactorily without the use of synchrotron
radiation. In particular, the superconduct-
ing wiggler beam line X17C provides a
high-energy beam with very high intensity
x-rays for probing tiny samples in dia-
mond-anvil cells. The infrared beam line
U4 provides high-intensity infrared radia-
tion for spectroscopic measurements of
samples in diamond-anvil cells, and is the

only facility of its type anywhere in the
World.

Another new development this year is
the installation of a Dilor micro-Raman
system by John Frantz and Bj om My sen for
examining the structures of silicate liquids
at high temperature. Heretofore, most
Raman studies of melts actually involved
measurements on glasses quenched from
high temperatures. Investigators were
forced to assume that the glasses were
representative of melts at temperature, but
there were many doubts about the validity
of this assumption. Now, Frantz and Mysen
have made extensive recordings of Raman
spectra on silicate melts at temperatures as
high as 1600°C and it appears that they
have opened up a new and exciting area of
melt research.

Douglas Rumble, Michael Palin, and
Thomas Hoering have developed a method
for laser fluorination of sulfide minerals
with fluorine gas. This technique allows
fast and precise in-situ micro-sampling on
three of the four sulfur isotopes, 32S, 33S,
and 34S, and will provide information on
mass transfer in sulfide system on a scale
that was previously inaccessible.

In addition to these initiatives related to
a new NSF Center and to new instrumenta-
tion, staff members, postdoctoral fellows
and associates, and visiting investigators
have been involved in a wide range of
research, ranging from racemization dat-
ing of ostrich shells to global convection of
the upper mantle. Much of this research is
described in brief summaries in this Annual
Report, thus continuing the Geophysical
Laboratory tradition of early communica-
tion of results before full papers are pub-
lished in scientific journals.

GEOPHYSICAL LABORATORY

Igneous and Metamorphic Petrology

Field studies

Global Convection and Hawaiian Upper
Mantle Structure

T. Neil Irvine

This report gives further development
of the global convection system proposed
by Irvine (1989). In this system, upper
mantle convection is stratified at both the
400- and the 670-km seismic
discontinuities, and the lower mantle fea-
tures an orthogonal framework of six prin-
cipal convection centers, or axes where
upwelling occurs beneath Iceland, Hawaii,
the Balleny Islands (near Antarctica), and
the Okavango delta (in Botswana), and
downwelling beneath Peru and the eastern
edge of Vietnam. In last year's Report
(Irvine, Annual Report 1989-1990, p. 3-
1 1), the concept of an upper mantle vortex
supercell was introduced for the Iceland
center and explored on the basis of seismic
data. Application of this concept is now
extended to the Hawaiian center.

Mantle Tomography and Hotspots

When the above global-convection sys-
tem was proposed, some support was cited
from seismic tomography, but the
tomographic maps then available portrayed
only broad features and frequently seemed
incompatible. More recent maps are not

completely consistent either (e.g., see com-
parisons made by Romano wicz, 1 99 1 ), but
a set by Inoue et al. (1990) is especially
interesting because the maps are unusually
detailed. Much of the detail correlates
meaningfully with surface geological fea-
tures (particularly volcanic hotspots), so
the data appear significant. Some features
are notably compatible with the mantle
convection relationships favored here.1

The Inoue et al. (1990) maps for nomi-
nal depths of 478-629 km and 1203-1435
km are illustrated in Fig. 1 , together with a
map by Woodhouse and Dzie wonski ( 1 989)
for 150 km. The relations for 1203-1435
km warrant particular attention because ( 1 )
they are dominated by two major zones of
low velocity, one under Hawaii, the other
beneath Okavango, and (2) the extensions
of these two zones in combination with
several small, low-velocity anomalies, en-
compass most of the world's hotspots. The
two major anomalies are both prominent in
several earlier tomographic maps (e.g.,
Giardini et al, 1987), but the velocity data
in these cases were smoothed to low-order
spherical harmonics, so the anomalies are
only broadly delimited. The anomaly loca-

1 In interpretations of mantle tomography here it
is assumed simply that seismically slow regions
are relatively warm and, thus, may represent zones
of convective upwelling, whereas fast regions are
cooler and, therefore, may reflect downwelling.
Compositional and phase differences and seismic
anisotropy are likely to be complicating factors
but cannot be considered here.

CARNEGIE INSTITUTION

150 km

Woodhouse and Dziewonski, 1989
Model U84L85/SH Z =1-8

+ + + +
+ + + +
+ + + +
+ + + +

-3%

478-629km

Inoue et al.,1990

::::: +

+ + + +
+ + + +
+ + + +
+ + + +

'////
VK

GEOPHYSICAL LABORATORY

tions on the Inoue et al. maps appear much
better resolved.

Neither the Iceland nor the Balleny cen-
ter is seismically slow at 1203-1435 km,
but Iceland is flanked by small low-veloc-
ity anomalies on the northwest and south-
east. Also, the Hawaii and Okavango
anomalies are beltlike at this depth and tend
to follow the great circle that includes Ice-
land and Balleny. In combination, the two
belts span almost half the Earth's circum-

ference. (A rather similar anomaly arrange-
ment is also indicated for 478-629 km in
Fig. IB.)

Several small anomalies in Fig. 1C do
not correlate with hotspots, and a few
hotspots are not associated with low ve-
locities (notably Tristan da Cunha, Fernando
de Norhana, and Martin Vas, all in a South
Atlantic region where relatively high ve-
locities are indicated); but at the present
state of the science, it seems more signifi-

1203-1435 km
Inoue et al., 1990

Volcanic hotspots

FIG. 1. Global tomography maps for nominal depths of (A) 150 km, redrawn from Woodhouse and
Dziewonski (1989) and (B and C) 478-629 km and 1203-1435 km, redrawn from Inoue etal. (1990). The
maps also show the global convection framework of Irvine (1 989).

The principal correlations between low-velocity anomalies and hotspots in Fig. 1C are as follows. The
Hawaiian anomaly spreads southward beneath the Samoa, Marquesas, Tahiti, and Austral McDonald
hotspots and comes close to Caroline (northeast of New Guinea). A strong small anomaly underlies the
southeastern Australia hotspot, and weaker ones match the Galapagos and San Felix (south of Peru), and
possibly Yellowstone. In the North Atlantic, a medium-sized anomaly is surrounded by the Azores,
Canary Islands, Madeira, Cape Verde Islands, and New England hotspots; and in the South Atlantic, a
strong small anomaly underlies the Cameroon hotspot, and a weak one matches St. Helena. Where the
Okavango anomaly extends into northern Africa, it and a small satellite are associated with four
continental hotspots; to the east and south, the Okavango anomaly encompasses the Comores, Reunion,
Marion Island, Crozet Islands, and Kerguelen hotspots.

CARNEGIE INSTITUTION

cant that many hotspots are matched. A
currently popular concept is that hotspots
are initiated by mantle plumes originating
in the D" seismic zone just above the core-
mantle boundary, and then continue to be
fed from this zone by thin, stemlike chan-
nels of up welling (e.g., Olson, 1990). Im-
pressive geoid and tomographic evidence
for a source region at the core -mantle bound-
ary has been available for some years (e.g.,
Chase, 1979; Crough and Jurdy, 1980;
Woodhouse and Dziewonski, 1989), but
the map in Fig. 1C contains what may well
be the first discernible geophysical indica-
tions of feeder stems at intermediate depths
in the lower mantle.

Stratified Mantle Convection

A prominent feature of the three maps in
Fig. 1 is that they are all very different; in
fact, a general observation from seismic
tomography is that the three mantle divi-
sions delimited by the 400- and 670-km
seismic discontinuities tend to exhibit con-
trasting relations. Although rarely cited,
this observation would seem a rather strong
argument in favor of at least some stratified
convection.

In the mantle convection relationships
proposed by Irvine (1989), it was assumed
that the 400-km interface is everywhere
mechanically coupled, and a combination
of thermal and mechanical coupling rela-
tionships was then devised for the 670-km
interface, designed to account for the gen-
eral tectonic features of the Earth's sur-
face. An unorthodox feature of the result-
ing arrangement is that, beneath the ob-

served zones of upwelling along the mid-
ocean ridges, there are zones of
downwelling in the depth interval 400-670
km. Some seismic evidence for this possi-
bility was cited from the tomographic maps
of Nataf et al. (1986), but it was acknowl-
edged to be equivocal. The relationships of
maps A and B of Fig. 1 suggest stronger
evidence for this possibility, although this
evidence too is not beyond question. In
particular, the maps indicate that, while the
regions beneath the mid-ocean ridges at
150 km (map A) are generally seismically
slow (as expected), those at 478-629 km
(map B) commonly exhibit relatively high
velocities. This contrast is especially con-
spicuous along the East Pacific Rise, where
it is most relevant to the Hawaiian convec-
tion relationships.

But, as explained by Inoue et al. ( 1 990),
their map for 478-629 km does not have as
high resolution as that for 1203-1425 km,
and they specifically stated that the East
Pacific Rise anomaly in the former is not
reliable. This does not mean that the
anomaly is invalid, however. Inoue et al.
noted also that an increase of seismic ve-
locity at 400 km beneath the East Pacific
Rise had been observed by Suetsugu and
Nakanishi ( 1 987) in a Rayleigh wave study;
and maps by Dziewonski and Woodhouse
(1987) of S-wave velocity variations at the
670-km discontinuity similarly show high
velocities beneath the Rise south of the
Galapagos. Thus, the possibility of sub-
ridge downwelling at 400-670 km, although
still wanting of strong support, is still con-
sidered viable.

GEOPHYSICAL LABORATORY

Hawaiian Relationships and Supercell
Structure

It is well established that the islands and
seamounts of the Hawaiian and Emperor
volcanic chains become progressively older
to the west and north from Hawaii. The
widely accepted explanation is that the
volcanoes formed in succession from a
relatively fixed mantle hotspot, currently
located beneath the Big Island, as the Pa-
cific plate drifted first northward, then
westward across it (cf. Clague and
Dalrymple, 1987). The recent volcanism
on Hawaii has been dominated by the tho-
leiitic eruptions of Kilauea andMauna Loa,
but the newest activity features eruptions
of alkalic basalt from the submarine vol-
cano Loihi on the south edge of the island.
Along the older parts of the volcanic sys-
tem, the seamounts at the Hawaiian-Em-
peror bend are 42-43 Ma in age; those at the
north end of the Emperor chain are 73-75
Ma.

In 1972, Jackson et al. pointed out that
the Hawaiian-Emperor volcanoes tend to
be paired, and portrayed them as being
distributed along an en echelon (discon-
tinuous) series of sigmoidal double lines.
Recently, Garcia et al. (1990) identified a
newly discovered submarine volcano
(named Makuhona) just west of Hawaii to
be the previously missing partner of the
volcano Kohala on the northern peninsula
of the island. With these observations as
background, I have attempted to pair the
volcanic structures of the entire system into
a set of more continuous double lines, as
shown in Fig. 2. En echelon overlap was
required through the Midway Islands, but

otherwise only two lines were necessary.
Inasmuch as the only control was small-
scale topography, the pairing is frequently
conjectural, but in an overall count, per-
haps 60 of some 70 possible pairs appear
credible. Thus, given that there are prob-
ably still other unrecognized eruption cen-
ters of the Makuhona type, the more con-
tinuous double lines seem realistic. The
contention is that they reflect the continu-
ous operation of the mantle convection
supercell outlined in parts B and C of Fig.
2 and explained below.

Figure 2 also shows the ocean-floor
troughs and arches that are associated with
the volcanic chains. The main Hawaiian
trough (here termed the "inner trough")
encircles the south side of the Big Island
and extends discontinuously westward on
both sides of the island chain-, with widths
locally exceeding 100 km. This trough has
long been ascribed to loading of the ocean
floor by the volcanic ridge (e.g., Moore,
1987). Outboard from it are broad arches,
or swells, 300-400 km wide. They are
topographically prominent only along the
younger half of the Hawaiian chain, where
their relief exceeds 1 000 m and brings them
to depths less than 5000 m (Fig. 2); but they
can be readily identified throughout, even
along the Emperor chain, by their distinc-
tive positive gravity signatures (see Haxby,
1987). Finally, fringing the swells around
the younger part of the Hawaiian chain is a
subtle "outer trough." This trough has
special importance in the present context
(see below).

Figure 3 depicts an Hawaiian upper
mantle convection supercell designed to
account for the paired volcano chains and

CARNEGIE INSTITUTION

FIG. 2. Maps of the Hawaiian and Emperor island and seamount chains in which the volcanic structures
are paired into two lines (four lines through the Midway Islands). Based on the topography map of
Mammerickx (1989), the ocean gravity map of Haxby (1987), and maps of earthquake epicenters and
young volcanoes from Moore (1987) and Garcia etal. (1990). See text for discussion of the upper mantle
convection supercell.

GEOPHYSICAL LABORATORY

WNW

EAST PACIFIC RISE

Mechanical coupling

Tholeiitic shield volcano
Early alkalic lavas
nner Trough
Outer trough

0 KM

PLAN VIEWS
Vortex core flow

0 down

11111*11 in i 400

FIG. 3. Schematic three-dimensional representation of the Hawaiian upper mantle vortex supercell. See
text for description.

related features.2 A principal feature of the
interpretation is the differential lateral flow
of the upper mantle layers. As with the
supercell proposed last year for Iceland,
the vorticity in the supercell is ascribed to
thermal tilting of this flow by heat rising
from the underlying lower mantle axis of
convective up welling (cf., Irvine, 1990,

2 Garcia et al. (1990) noted, on the basis of an
experimental study of mantle plume dynamics by
Richards and Griffiths (1989), that a convection
structure of this general type might account for the
pairing of the Hawaiian volcanoes; and Sleep
(1990) used a "stagnation streamline" model akin
to part of the structure in Fig. 3 to estimate buoy-
ancy flux values for the Hawaiian and other mantle
plumes.

Fig. 1). The differential flow also leads to
the downwelling at 400-670 km under the
East Pacific Rise, as discussed in relation to
Fig. 2B.

In the envisaged action of the supercell,
a swath of the lower part of the lithosphere
is stripped away (delaminated) and replaced
by asthenospheric material upwelling from
near the 400-km interface. The stripping
process depresses the outer ocean-floor
trough, and the buoyancy of the replace-
ment material elevates the topographic
swells. The swells eventually subside as
the replacement material gradually cools
and contracts, but through its increased

CARNEGIE INSTITUTION

density, they continue to have their distinc-
tive gravity signatures. Further postulates
are (1) that the double -vortex circulation
holds the supercell in place as a standing
vortex against the laterally flowing mantle
layers (see Irvine, 1990, Fig. 1), (2) that
magma melting occurs mainly through the
adiabatic rise of fertile peridotite along the
two vortex axes of upwelling, and (3) that
these two axes also deliver the magma to
paired release points at the base of the
lithosphere, from where it rises to the paired
volcanoes at the surface.

The supercell structure can also be used
to rationalize much of the general history
of individual Hawaiian volcanoes (for sum-
mary, see Clague and Dalrymple, 1987).
By concept, the magma from the main
release points, having melted at relatively
high pressures, is picritic tholeiite in com-
position, and its eruption produces large
shield volcanoes like Mauna Loa and
Kilauea. But it can be postulated too that
some alkalic magma should form in the
surrounding, spreading parts of the vortex
cells by differentiation of trapped intersti-
tial melt at the more moderate pressures of
this environment. This magma could be
released both in advance of the tholeiitic
shields, as at Loihi, and in arrears of them,
as in the "alkalic post-shield stage" ob-
served in volcanoes such as Mauna Kea
and Hualalai. I suggest too that, through its
buoyancy, this alkalic magma may tend to
accumulate in substantial quantities at the
tops of the inferred vortex axes of
downwelling, from where it might be
erupted at a considerably later stage in the
history of a volcano if the volcano hap-
pened to pass across such an axis by virtue

of the plate motion. This kind of late
eruption could correspond to the "alkalic
rejuvenated stage" that is observed in most
of the Hawaiian volcanoes between Maui
and Necker Island.

A fundamental general tenet in this
analysis is that magma from the mantle is
primarily released to the lithosphere from
specific, critical flow points in the
asthenospheric convection system.

References

Chase, C. G., Subduction, the geoid, and lower
mantle convection: Nature, 282, 462-468,
1979.

Clague, D. A., and G. B. Dalrymple, The Hawai-
ian-Emperor volcanic chain: Part I. Geologic
evolution, U. S. Geol. Surv. Prof. Paper 1350,
5-54, 1987.

Crough, S. T , and D. M. Jurdy, Subducted litho-
sphere, hotspots, and the geoid, Earth Planet.
Sci. Letters, 48, 15-22. 1980.

Dziewonski, A. M., andWoodhouse, J. H., Global
images of the Earth's interior, Science, 236,
37-48, 1987.

Garcia, M. O., M. D. Kurz, and D. W. Muenow,
Mahukona: The missing Hawaiian volcano,
Geology, 18, 1111-1114, 1990.

Giardini, D., L. Xiang-Dong, and D. H.
Woodhouse, Three-dimensional structure of
the Earth from splitting in free-oscillation spec-
tra, Nature, 325, 405-41 1, 1987.

Haxby, W. F., Map of the gravity field of the
world's oceans, Natl. Oceanic Atmos. Adm.
Rpt. MGG-3, 1987.

Inoue, H., Y. Fukao, K. Tanabe, and Y. Ogata,
Whole mantle P- wave travel time tomography,
Phys. Earth Planet. Interiors, 59, 294-328,
1990.

Irvine, T. N., A global convection framework:
concepts of symmetry, stratification and sys-
tem in the Earth's dynamic structure, Econ.
Geol, 84, 2059-21 14, 1989.

Jackson, E. D., E. I. Silver, and G. B. Dalrymple,
Hawaiian-Emperor chain and its relation to
Cenozoic Circum-Pacific tectonics, Geol. Soc.
Amer. Bull., 83, 601-618, 1972.

Mammerickx, J., Bathymetry of the North Pacific
Ocean, Geol. Soc. Amer., The Geology of
North America, N, 1989.

GEOPHYSICAL LABORATORY

Moore, J. G., Subsidence of the Hawaiian Ridge,
U. S. Geol. Surv. Prof. Paper 1350, 85-100,
1987.

Nataf, H.-C, I. Nakanishi, and D. L. Anderson,
Measurements of mantle wave velocities and
inversion for lateral heterogeneity and anisot-
ropy: III, Inversion, /. Geophys. Res., 91,
7261-7308, 1986.

Olson, P., Hot spots, swells and mantle plumes, in
M. P. Ryan, Magma Transport and Storage,
New York, J. Wiley & Sons, 33-51, 1990.

Romano wicz, B., Seismic tomography of the
Earth's mantle, Ann. Rev. Earth Planet. Sci.,
79,77-99, 1991.

Sleep, N., Hotspots and mantle plumes: some
phenomenology, /. Geophys. Res., 95, 6715-
6736, 1990.

Suetsugu,D.,andI.Nakashini, Three-dimensional,
velocity map of the upper mantle beneath the
Pacific Ocean as determined from Rayleigh
wave dispersion, Phys. Earth Planet. Interi-
ors, 47, 205-229, 1987.

Woodhouse, J. H., and A. M. Dziewonski, Seis-
mic modelling of the Earth's large scale three-
dimensional structure, Phil. Trans. R. Soc.
Lond., A 328, 291-308, 1989.

Richards, M. A., and R. W. Griffiths, Thermal
entrainment by depleted mantle plumes, Na-
ture, 342, 900-902, 1989.

Megacrystalline Dunites and

Peridotites:
Hosts for Siberian Diamonds

N. P. Pokhilenko* , D. G. Pearson ** ,
F. R. Boyd, andN. V. Sobolev*

Several investigations have identified
xenoliths, consisting primarily of
ultracoarse crystals of olivine, which ap-
pear to be fragments of the principal host
rocks of Siberian diamonds (Sobolev, 1974;
Pokhilenko et ai, 1977; Soboley et al.,
1984). Twenty -three of these xenoliths

Inst, of Mineralogy & Petrology, Siberian Branch
of the USSR Academy of Sciences, Novosibirsk
Dept. of Terrestrial Magnetism, Carnegie Instn.
of Washington, Washington, D. C. 20015.

contain diamonds, and many more contain
garnets with a compositional range very
similar to the range for garnets included in
diamonds. These megacrystalline rocks
have been found as xenoliths in the
kimberlites of the Daldyn-Alakit region
and in some pipes of the Upper Muna
group. They are especially abundant in the
xenolith-rich Udachnaya kimberlite, site
of one of the world's richest diamond mines.

The Siberian megacrystalline rocks dif-
fer from diamondiferous peridotites from
southern Africa, the only other region in
which a number of such xenoliths have
been found. African occurrences are pri-
marily lherzolites and harzburgites having
garnets and associated minerals with com-
positional ranges that depart significantly
from the majority of diamond inclusions
(Viljoen etal., 1991; Boyd etal., in prepa-
ration). Thus the principal host rocks for
African diamonds have not yet been dis-
covered as articulated xenoliths.

The relative abundances and composi-
tional ranges of the olivine and associated
minerals indicate that most of the Siberian
megacrystalline rocks are extremely re-
fractor)'. They may be residues or cumu-
lates of melting events in which a large
proportion of melt was removed. Olivine
forms more than 90% of most specimens.
The primary modal proportions are diffi-
cult to estimate because of disaggregation
during eruption, but it is likely that these
ultracoarse rocks were olivine-rich. The
molar Mg/(Mg+Fe) for the olivine is pre-
dominantly 0.92-0.95, and most of the gar-
nets are strongly subcalcic and Cr-rich (Fig.
4). Mineral assemblages of the 33 new

CARNEGIE INSTITUTION

Table 1. Mineral assemblages in megacrystalline
olivine-rich xenoliths from the Udachnaya
kimberlite, U.S.S.R.

Assemblage

No. specimens

olv + gar

olv + gar + chr

olv + gar + diam

olv + gar + cpx

olv + gar + opx

olv + chr

olv + opx + chr

olv + gar + chr + diam

olv + gar + opx + cpx + chr

specimens chosen for this investigation
have a predominance of olivine + garnet
(Table 1). Five are diamond-bearing.

Isotope Results

Preliminary neodymium isotope analy-
sis of two hand-picked, acid-washed garnet
separates (Uv 70/76 and Uv 49/76) reveal
that the megacrystalline rocks, having ex-
perienced an ancient trace element enrich-
ment event, show similar Nd isotopic sig-
natures to peridotite suite garnet inclusions
in diamond from southern Africa analyzed
by Richardson et al. (1984). The two
samples record £Nd values of -28 to -33 at
350 Ma (the time of pipe emplacement),
and yield model ages relative to the Bulk
Earth reservoir (CHUR) of 2.0 to 2.7 Ga.
(Fig. 5). Although the latter age is late
Archaean, it is substantially younger than
the 3.3 Ga. model ages obtained for the
diamond inclusions by Richardson et al.
(1984). The younger model ages and the
age-range obtained from the two samples
from Udachnaya may be attributed to the
more "open system" behavior of the small,
coarse-grained megacrystalline samples

5 8

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Cr203,wt%

Fig. 4. A plot of CaO against O2O3 for garnets
from megaciystalline xenoliths of different asso-
ciations: 1 - without chromite; 2 - with chromite;
3 - diamond-bearing without chromite; 4 - dia-
mond-bearing with chromite. Data from this
study and Sobolev et al. (1984).

(compared with the armored diamond in-
clusions), which could have allowed the
xenolith garnets to be subjected to later
interaction with a lighter rare-earth-enriched
component. Further analyses of a suite of
larger garnets from the megacrystalline
rocks are being undertaken to resolve this
problem. However, the results obtained
indicate that the base of both the Kaapvaal

0.512

0 1000 2000 3000 4000

Time, Ma

Fig. 5. Neodymium isotope evolution diagram for
two-garnet separates from the Udachnaya
megacrystalline. Vertical arrows indicate model
ages relative to the Bulk Earth or CHUR evolution
curve. Diamond symbols represent the isotopic
composition of the samples at the time of pipe
emplacement 350 Ma ago.

GEOPHYSICAL LABORATORY

and Siberian cratons experienced similar
incompatible element enrichment events
very early in their evolution.

Petrography

The megacrysts are friable and are not
easily separated intact from hard kimberlite.
The largest has a long dimension of 19 cm,
but most specimens are ovoidal with di-
mensions of the order of 5 cm. Olivine
crystals range up to 10 cm, and in a few
specimens smaller grains of olivine (6-9
mm) with differing optic orientation are
interspersed. Prominent parting, resem-
bling cleavage, is characteristic. Rounded
garnets are much finer-grained, 0.2-6 mm,
and form up to 5 modal percent. The
predominant low-Ca garnets are lilac -purple
but those that are rich in Ca as well as Cr are
dark purple or gray or even green. Alter-
ation of garnet to kelyphite is variable;
some grains with minor kelyphite are
euhedral with a rhombic dodecahedral
morphology complicated by uneven devel-
opment of faces. The distribution of gar-
nets is relatively regular in the larger xeno-
liths.

Enstatite is present in about 20% of the
megacrystalline rocks, but clinopyroxene
is much less common. Octahedra of chro-
mite, 0.1-2 mm, form less than 1 modal
percent, and some of them appear to show
signs of resorption.

Crystals of diamond in these rocks are
characteristically clear, shaip-edged octa-
hedra, some with spinel twins, and range up
to 4 mm. Colored diamonds and those that
are corroded or cracked are unusual. Small

plates of graphite (0.2-1.5 mm) have also
been observed both with and without dia-
mond. The abundances of diamond and
graphite are usually insignificant but one
olivine-pyrope specimen having a volume
of only 2.5 cm3 contains four diamonds
together with cavities from which two ad-
ditional diamonds have broken away.

Mineral Chemistry

Most olivines in the megacrystalline
rocks have Mg/(Mg+Fe) in the range 0.92
-0.95, but two olivine-rich wehrlites in
which the garnets have both high Ca and
high Cr (Fig. 4) contain substantially more
Fe-rich olivines, 0.87 - 0.89 (Fig. 6). Sur-
prisingly, a variation in NiO from 0.32 to
0.40 does not correlate with mg number.
Values for CaO and Cr203 in the olivines
do not exceed 0.06 wt %.

Garnets in the megacrystalline xeno-
liths can be assigned to three parageneses
on the basis of their CaO and Cr203 con-
tents: proportions for more than 200 speci-
mens thus far studied (this Report and
Sobolev et al., 1984) are harzburgite-dun-
ite (80%), lherzolite (15%), and wehrlite
(5%). Concentrations of FeO andTi02 are
least in the harzburgite-dunite group (Tables
2 and 3). There is a positive correlation
between Ti02 and CaO in these garnets,
apparently reflecting degree of depletion
(Fig. 6B). The chromites are predomi-
nantly rich in Cr203 (58-65 wt %) but
spinels with Cr203 as low as 21.9 wt %
have been found. A good correlation in Cr/
(Cr + Al) between the garnets and chro-
mites (Fig. 6C) is evidence of equilibra-

CARNEGIE INSTITUTION

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CARNEGIE INSTITUTION

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20 40 60 80

Cr/(Cr+AI), %, chr

100

Fig. 6. Compositional variations in minerals of the
megacrystalline rocks. Symbols as in Fig. 4. a) a
plot of NiO against Mg number for olivines of
megacrystalline rocks, b) a plot of CaO against
Ti02 for garnets from megacrystalline rocks, c) a
plot of Cr/(Cr + Al) ratio for garnets against Cr/(Cr
+ Al) ratio for chromites from megacrystalline
rocks.

700

30 40 50 60

Pressure, kbar

Fig. 7. Temperature-Pressure estimates for four
xenoliths of megacrystalline rocks from
Udachnaya pipe having coexisting pyrope and
enstatite. Temperature was calculated with the
olivine-garnet thermometer (O'Neill and Wood,
1979) and pressure from the isopleths of
MacGregor(1974).

tion. Enstatites have low AI2O3 (Table 2),
reflecting relatively low temperatures and
high pressures of crystallization.

Thermobawmetry

The temperature of equilibration of a
xenolith can be estimated from the parti-
tion of Fe and Mg between olivine and
garnet (O'Neill and Wood, 1979), but
enstatite must also be present to obtain an
estimate of the pressure or depth of equili-
bration. Four of the megacrystalline rocks
contain enstatite as well as olivine and
garnet and<F-r estimates for them (Fig. 7)
plot close to or within the diamond stabil-
ity field.

Puzzling features of the P-Tplot in Fig.
7, however, are that the estimated points for
megacrystalline rocks show a wide disper-
sion in pressure and that they plot in a
temperature range below the shield
geotherm (40 mWm^) and below estimates

GEOPHYSICAL LABORATORY

made for garnet lherzolite xenoliths from
Udachnaya (Boyd, 1984). One of the
megacrystalline rocks contains diopside as
well as olivine and garnet (Uv-624/86, Table
2), and its temperature can be estimated
with the diopside solvus, giving 800°C, as
well as with the Gar/Olv thermometer,
which gives 840°C. The approximate agree-
ment suggests that the discrepancy with the
lherzolite data may not be a failure of
thermobarometry. The dispersion in pres-
sure for the megacrystalline rocks could be
evidence of a wide range in depth of equili-
bration. More data are clearly needed,
however, before these questions can be
properly addressed.

Experimental studies of coexisting solid
solutions of spinel and garnet show that the
solubility of the knorringite component
(Mg3Cr2Si30i2) in pyrope increases with
increasing temperature and pressure
(Malinovsky andDoroshev, 1975). Pyrope
garnet that is rich in Cr can coexist with

03
O

4 6 8 10

O2O3 ,Wt.%

12 14

Fig. 8. A plot of CaO against Q2O3 with isopleths
of knorringite component in garnet estimated from
experimental results (Malinovsky and Doroshev,
1975). Fields for graphite and diamond are calcu-
lated from these experimental data and the esti-
mated temperature and pressure of the diamond-
graphite transition in the Siberian mantle.

chromite only at pressures within the dia-
mond stability field (Kesson and Ringwood,
1989). The experimental data can be used
to estimate the compositions of garnets in
equilibrium with Cr-spinel at P-T condi-
tions corresponding to the diamond-graph-
ite transition in the Siberian lithosphere
(Fig. 8). The compositions of garnets in
megacrystalline rocks that contain diamond
and chromite are consistent with the field
for diamond shown in Fig. 8 and with the
knorringite isopleths based on experiment.
Analyses for garnets in graphite + garnet +
chromite assemblages (not plotted in Fig.
8) all have less than 15% knorringite and
are thus also consistent (Pokhilenko et al.,
1988). Thus, it may be possible to develop
a useful barometer based on the Ca and Cr
contents of the garnets for assemblages that
include both pyrope and chromite.

References

Kesson, S. E., and A. E. Ringwood, Slab-mantle
interactions 2. The formation of diamonds,
Chem. Geol., 78, 97-118, 1989.

Malinovsky, I. Yu., and A. M. Doroshev, Stability
of garnet of pyrope-knorringite row in the field
7=1000-1500°C and P=20-50 kb, Nauka.
Novosibirsk, p. 23-31, 1975.

O'Neill, H. St. C, and B. J. Wood, An experimen-
tal study of Fe-Mg partitioning between garnet
and olivine and its calibration as a
geothermometer, Contrib. Mineral. Petrol. 70,
59-70, 1979.

Pokhilenko, N. P., A. S. Rodionov, T. M. Blinchik,
and E. V. Malygina, Graphite-diamond phase
transition and its significance for estimations of
P-T conditions of equilibrium of ultrabasic
xenoliths, Ext. Abstr. vol. Intern. Symposium
on Composition and Processes of Deep-seated
Zones of Continental Lithosphere, pp. 64-65,
Novosibirsk, 1988.

Pokhilenko, N. P., N. V. Sobolev, and Yu. G.
Lavrent'ev, Xenoliths of diamondiferous ultra-
mafic rocks from Yakutian kimberlites, 2nd
Int. Kimb. Conf. Ext. Abstr. Vol. Santa Fe,
1977.

CARNEGIE INSTITUTION

Richardson, S. H., J. J. Gurney, A. J. Erlank, and
J. W. Harris, Origin of diamonds in old en-
riched mantle, Nature, 310, 198-202, 1984.

Sobolev, N. V., Deep Seated Inclusions in
Kimberlites and Problem of Upper Mantle
Composition, p. 1-264 Nauka, Novosibirsk,
1974 (English Translation), AGU, 1977.

Sobolev, N. V., N. P. Pokhilenko, and E. S.
Yefimova, Diamond-bearing peridotite xeno-
liths in kimberlites and the problem of the
origin of diamonds, Sov. Geol. Geophys., 25,
62-76, 1984.

Viljoen, K. S., D. H. Robinson, and P. M. Swash,
Diamond and graphite peridotite xenoliths from
the Roberts Victor Mine, 5th Intern. Kimb.
Conf. Araxa, Brazil, Ext. Abstr. Vol., 1991.

Mantle Metasomatism: Evidence from a

MARIE) - Harzburgite Compound

Xenolith

F. R. Boyd

Interpretation of metasomatized rocks
in the Earth's crust depends critically on the
study of outcrops. We are not fortunate,
however, in having exposures of the deep
portions of continental cratons which have
been the sites of metasomatic events over
the course of 3-4 billion years. In the
absence of outcrops, the rare discoveries of
xenoliths exhibiting contacts between rock
types have particular importance. This is
especially true if the rocks are fragments of
an igneous intrusion and a metasomatized
conduit wall.

A xenolith containing a contact between
a mica-rich igneous cumulate and a
metasomatized peridotite was discovered
and made available for study by staff of the
Anglo Axnerican Research Laboratories.
The xenolith was recovered from the com-
bined coarse concentrate of the De Beers
mines in Kimberley.

Mica-rich xenoliths from the Kimberley

pipes include varieties of both igneous and
metasomatic origin. Those believed to be
igneous have been designated by the acro-
nym MARID, based on the names of the
constituent primary minerals: mica, am-
phibole, rutile, ilmenite, and diopside
(Dawson and Smith, 1977). Metasomatized
peridotites contain a variety of introduced
phases that include phlogopite, potassic
richterite and pargasite amphiboles, il-
menite, rutile, and a number of exotic po-
tassium and barium titanates (Erlank et al.,
1987). It has been a problem to understand
the nature of the metasomatizing fluids and
whether or not they are related to kimberlite.
Is kimberlite the magma from which
MARID rocks are derived and are MARID
rocks the remnants of metasomatic sources?
Only one xenolith containing a contact
between a MARID rock and a peridotite
has previously been described (Waters et
al., 1989). This xenolith was interpreted as
a fragment of an intrusive with attached
wall rock and study of it has helped to
establish the hypothesis that MARID rocks
are remnants of sources of metasomatic
fluids. Extensive post-metasomatic alter-
ation of the peridotite to carbonate in this
xenolith, however, has obscured the pri-
mary mineralogy and the metasomatic im-
print.

Petrography

The sawed fragment of the compound
xenolith analyzed in the present study is 6
cm in maximum dimension and 2 cm thick;
the fragment is numbered FRB 1455. The
MARID portion of the xenolith consists of

GEOPHYSICAL LABORATORY

over 95 % phlogopite, predominantly in the
form of tablets, 1-2 mm in length, aligned
parallel to the contact with the peridotite.
The coarse mica is strained, having undu-
late extinction and bent cleavage. Finer-
grained mica neoblasts are interspersed
between the coarse tablets. The coarse
mica is pale, but the neoblasts and mantles
on the mica tablets have a darker, yellow-
ish-brown color.

Granules of diopside a few tenths of a
mm in diameter are dispersed between the
mica tablets and clustered with serpentine
in lenticles. Elongate grains of rutile rang-
ing up to a mm are less abundant. Small
patches of calcite that may be primary are
widely dispersed and enclose euhedral crys-
tals of sphene, 0.1-0.2 mm. Sphene is an
unusual mineral in mantle rocks, but
Dawson and Smith (1977) have noted its
occurrence as an accessory phase in a
MARID xenolith from the Wesselton mine.

The harzburgite consists primarily of
coarse olivine, ranging up to 1 cm, with a
minor proportion of smaller grains of
enstatite. Olivine neoblasts locally form
interstitial zones but are insufficiently abun-
dant to envelope the primary grains. Coarse
crystals of phlogopite, mantled and seamed
with finer-grained mica, range up to 8 mm
and have a color and degree of strain simi-
lar to the mica in the adjacent MARID rock.
The section analyzed contains several 1-
mm grains of garnet enveloped in fine-
grained mica, forming a replacement tex-
ture like that cited by Erlank et al. (1987)
and observed in Jagersfontein peridotites
(Boyd and Mertzman, 1987). The
harzburgite also contains irregularly-
shaped, coarse blebs of ilmenite mantled

by rutile which are up to a centimeter in
maximum dimension. The contact between
the peridotite and MARID is sharp; minor
serpentine and ilmenite lenticles are local-
ized in the peridotite adjacent to the con-
tact.

Mineral Chemistry

The mica and accessory phases in the
MARID portion of xenolith 1455 have
compositions that are comparable to those
previously reported for MARID rocks. The
phlogopite and diopside have low Cr203
(Table 4); the diopside is also relatively low
in AI2O3. The mg number of the coarse
phlogopite is 0.87, near the high end of the
range reported by Dawson and Smith
(1977). The diopside with an mg number
of 0.86 is markedly more Fe-rich than the
Cr-diopsides in common garnet lherzolites.
The darker mica that forms neoblasts and
mantles on the coarse tablets is enriched in
Ti but not in total Fe.

The olivine and enstatite in the associ-
ated peridotite have extremely variable Mg/
(Mg+Fe). Cores of large olivine crystals
are Mg-rich (0.94) but neoblasts and mar-
gins on the coarse grains are as Fe-rich as
0.88 (Tables 4, 5). These variations are
irregular, differing widely from grain to
grain; they are not systematic in reference
to the MARID-peridotite contact. Enstatite
grains have comparable enrichment in Fe
and the secondary enstatite has higher Ca
and Ti. Enstatite close to the MARID
contact, however, is markedly depleted in
Al (Table 5). The pale mica and darker
mica in the harzburgite have compositions

CARNEGIE INSTITUTION

Table 4. Compositions of primary minerals in MARID
1455, Kimberley, RSA.

harzburgite compound xenolith, FRB

Harzburgite

MARK)

Olivine

Enstatite

Garnet

Mica

Diopside

Rutile

Sphene

Si02

41.5

57.3

41.4

41.9

54.3

<0.03

30.4

Ti02

<0.03

0.10

0.70

1.77

0.40

97.2

41.9

AI2O3

<0.03

0.68

20.5

10.1

0.50

0.94

0.55

Cr203

0.04

0.48

3.26

0.18

0.48

1.89

0.03

FeO

5.74

4.63

8.20

6.22

5.11

0.09

0.79

MnO

0.07

0.11

0.39

0.04

0.13

<0.03

0.04

MgO

53.1

36.4

20.6

23.5

17.6

0.09

0.10

CaO

<0.03

0.50

5.51

<0.03

20.6

0.09

24.8

Na20

n.d.

0.21

0.11

0.10

0.94

n.d.

2.50

K20

n.d.

10.4

<0.03

n.d.

0.07

NiO

0.36

0.08

<0.03

0.15

0.03

0.08

total

100.8

100.5

100.7

94.4

100.09

100.3

101.3

Mg/(Mg+Fe)

0.943

0.934

0.818

0.871

0.861

Table 5. Compositions of secondary minerals occurring as discrete grains or mantles on
primary grains in MARID-harzburgite compound xenolith, FRB 1455, Kimberley, RSA.

Harzburgite

MARID

Olivine

Enstatite

Ilmenite

Rutile

Mica

Si02

40.5

56.4

<0.03

40.6

40.0

Ti02

0.03

0.30

53.6

94.0

3.73

3.64

A1203

<0.03

0.22

0.81

0.93

11.1

Cr203

0.04

0.10

2.59

2.97

0.19

0.23

FeO

11.5

8.16

30.4

0.29

5.29

5.91

MnO

0.13

0.27

<0.03

0.08

0.07

MgO

48.9

33.9

13.6

0.10

23.2

22.7

CaO

0.06

0.91

0.04

<0.03

Na20

n.d.

0.14

n.d.

0.20

0.17

K20

n.d.

10.2

10.3

NiO

0.38

0.07

0.19

<0.03

0.12

total

101.5

100.3

101.5

98.3

94.7

94.2

Mg/(Mg+Fe)

0.884

0.881

0.443

0.887

0.873

very similar to counterparts in the MARID
portion of the xenolith. Rutile in both the
MARID and peridotite contains several
percent of Cr203, a distinguishing feature
noted for MARID rutiles by Dawson and
Smith (1977).

Thermobarometry

The occurrence of olivine and garnet
together with enstatite makes it possible to
estimate the temperature and depth of equili-
bration of the harzburgite- MARID assem-

GEOPHYSICAL LABORATORY

Table 6. Estimates of the temperature and depth of equilibration of the harzburgite-
MARID xenolith, FRB 1455.

Thermometer

Barometer Temperature,

°C

Depth, km

olv-gar-opx

O'Neill-Wood (1979)
O'Neill-Wood (1979)

MacGregor(1974)
Nickel-Green (1985)

opx-gar-opx

530
560

50
70

Harley (1984)

MacGregor(1974)

850

115

Discussion

blage. Such estimates have not previously
been possible for MARK) xenoliths be-
cause they do not in themselves contain a
necessary combination of phases. Enrich-
ment of some of the olivine and enstatite in
Fe creates an obvious uncertainty, but the
most magnesian compositions are taken to
be primary (Table 4). This assumption may
not be correct for the enstatite, however,
because it has a slightly lower mg number
than the olivine, the reverse of the usual
relationship. Combining the olivine-gar-
net and the orthopyroxene-garnet thermom-
eters with two Al-enstatite barometers pro-
vides results that suggest a shallow mantle
origin for this xenolith (Table 6), perhaps
close to the upper limit of garnet peridotite
stability. There is a discrepancy of about
300°C, however, between temperatures cal-
culated with the olivine-garnet and enstatite-
gamet thermometers, and because of the
temperature dependance of the barometer
there is a corresponding discrepancy in
depth. This is probably due to slight en-
richment of the primary enstatite in Fe, and
the olivine-gamet estimates are believed to
be more reliable.

The pronounced gradients in Fe/Mg in
the olivine of xenolith 1455 may be evi-
dence that the metasomatism occurred in
association with the eruptions that pro-
duced the Kimberley diatremes. Such gra-
dients are unlikely to persist in the upper
mantle for periods of many million years.
The eruption of the six large diatremes at
Kimberley, each with multiple units of
kimberl'te, was a complex volcanic event,
however, and the origin of the xenolith may
have involved more than one body of
magma.

A second factor of importance is the
relatively low ambient temperature and
depth that are reflected by the composi-
tions of the primary minerals in the
harzburgite (Table 6). It appears that the
metasomatic process occurred at shallow
levels in the craton, far removed from the
top of the asthenosphere and from the zone
of kimberlite generation. In such circum-
stances it is probable that the
metasomatizing agent was either the
MARID magma or the kimberlite itself,
and possibly they are the same. A relation-

CARNEGIE INSTITUTION

ship between MARID rocks and kimberlite
has previously been proposed (Dawson
and Smith, 1977). The only alternative
source for the MARID rock that is known
to exist are intrusions of mica-rich (Group
II) kimberlite at nearby Loxtondal that have
an age 40 My older than the Kimberley
diatremes (E.M.W. Skinner, personal com-
munication).

Many of the peridotite and dunite xeno-
liths from Kimberley have been subject to
Fe-Ti metasomatism with introduction of
ilmenite and development of Fe-enriched
marginal zones on primary olivine and py-
roxene. These effects are believed to have
been caused by kimberlite magmatism
(Boyd etal., 1983), and they have appeared
to be distinct from the introduction of coarse,
Mg-rich mica and less commonly pargasite
that may have originated in events that long
pre- dated kimberlite eruption. Crystalliza-
tion of metasomatic K-richterite in perido-
tite, however, has involved introduction of
mica together with titaniferous phases
(Erlankeftf/., 1987). The latter metasomites
appear to be linked to MARID intrusions
by similarities in mineralogy and by the
evidence from xenolith 1455 and from the
similar compound xenolith studied by
Waters etal. (1989). Some of the variations
in metasomatic imprint may be conse-
quences of the varying state of kimberlite-
related metasomatizing fluid and varying
depth at which the reactions occurred.

Compositional changes induced in the
peridotite of xenolith 1455 are extremely
irregular, and their magnitude does not
vary systematically with distance from the
MARID contact on a scale of several cen-

timeters. Moreover, the irregular spatial
variation does not seem to reflect the origi-
nal presence of some other contact. Most
likely, the agent of metasomatism was a
hydrothermal fluid derived from nearby
magma that penetrated the harzburgite ir-
regularly along fractures and grain bound-
aries.

It is suggested that the MARID rock of
xenolith 1455 is an aggregate of mica with
minor diopside and rutile that was plated on
the peridotite wall of a dike or other conduit
through which kimberlite (?) magma was
erupting. The magma from which MARID
rocks have crystallized has been proposed
to be lamproite (Waters, 1987). There are
no lamproite s in the Kimberley area, how-
ever, and kimberlite itself is an alternative
possibility. The metasomatism could have
been produced by the same volume of melt,
but an earlier or later origin during the
restricted period of kimberlite magmatism
seems possible. The more darkly colored
Ti-rich mantles on the mica in both the
MARID and peridotite portions of the xe-
nolith may be evidence of a change in
composition of the metasomatic fluid or
even a two-stage process.

References

Boyd, F. R., and S. A. Mertzman, Composition
and structure of the Kaapvaal lithosphere,
southern Africa, in Magmatic Processes:
Physicochemical Principles, B. O. My sen,
ed., Geochemical Society Special Publication,
No. l,p. 13-24,1987.

Boyd, F. R., R. A. Jones, and P. H. Nixon, Mantle
metasomatism: the Kimberley dunites,
Carnegie Instn. Washington Year Book, 82,
330-336, 1983.

GEOPHYSICAL LABORATORY

Dawson, J. B., and J. V. Smith, The MARID
(mica-amphibole-rutile-ilmenite-diopside)
suite of xenoliths in kimberlite, Geochim.
Cosmochim. Acta, 41, 309-323, 1977.

Erlank, A. J., F. G. Waters, C. J. Hawkes worth, S.
E. Haggerty, H. L. Allsopp, R. S. Rickard, and
M. Menzies, Evidence for mantle metasoma-
tism in peridotite nodules from the Kimberley
Pipes, South Africa, in Mantle Metasomatism,
M. Menzies and C. J. Hawkesworth, eds.,
Academic Press, 1987.

Harley, S. L., An experimental study of the parti-
tioning of Fe and Mg between garnet and
orthopyroxene, Contrib. Mineral. Petrol, 86,
359-373, 1984.

MacGregor, J. D., The system MgO-Al203-SiC>2:
Solubility of AI2O3 in enstatite for spinel and
garnet peridotite compositions, Amer. Min-
eral., 59, 110-119, 1974.

Nickel, K. G., and D. H. Green, Empirical
geothermobarometry for garnet peridotites and
implications for the nature of the lithosphere,
kimberlite s and diamonds, Earth Planet. Sci.
Lett., 73, 158-170, 1985.

O'Neill, H. St. C., andB. J. Wood, An experimen-
tal study of Fe-Mg partitioning between garnet
and olivine and its calibration as a
geothermometer, Contrib. Mineral. Petrol.,
70, 59-70, 1979.

Waters, F. G., A suggested origin of MARID
xenoliths in kimberlites by high pressure crys-
tallization of an ultrapotassic rock such as
lamproite, Contrib. Mineral. Petrol., 95, 523-
533, 1987.

Waters, F. G., A. J. Erlank, and L. R. M. Daniels,
Contact relationships between MARID rock
and metasomatized peridotite in a kimberlite
xenolith, G eo chemical J ., 23, 11-17, 1989.

Boron and Beryllium Concentrations in

Subduction-Related Metamorphic Rocks

of the catalina schist! implications for

subduction-zone recycling

Gray E. Bebout, Jeffrey G. Ryan* and
William P. Leeman**

Enrichments in B and 10Be in arc volca-
nic rocks have been interpreted to reflect
slab additions to arc source regions via
hydrous fluids or silicate liquids (Morris et
ai, 1990; Ryan and Langmuir, 1991). The
two-component mixing relations among B,
10Be, and 9Be demonstrated by Morris et al.
(1990) are believed to result from the addi-
tion of a homogenized slab-derived compo-
nent, probably a hydrous fluid which
strongly fractionates B from Be. The con-
centrations of B and Be in sediments and
crustal materials believed to be subducted
are reasonably well characterized. Exten-
sive B and Be data sets exist for arc volcanic
rocks (Teraef a/., 1986; Morris etal, 1990;
Ryan and Langmuir, 1988, 1 991; Leeman et
al., 1990); thus, the volcanic output of these
elements can be estimated (see Ryan and
Langmuir, 1991, for discussion of B). How-
ever, relatively little is known about the
effects of slab metamorphism on the concen-
trations of 3, Be, and other trace elements in
subducted rocks. These metamorphic pro-
cesses may dictate the efficiency with which

these elements are transferred from surface
reservoirs to arc source regions; Moran et

* Department of Terrestrial Magnetism, Carnegie
Institution of Washington
** Keith -Wiess Geological Laboratories, Rice
University, Houston, Texas 7725 1

CARNEGIE INSTITUTION

al. (1991) discuss the consequences of B
loss due to slab metamorphism for the
mass -balance of B inputs to arc source
regions. Also of particular interest are the
relative capabilities of metamorphic
de volatilization and migmatization pro-
cesses to fractionate B and Be sufficiently
to produce the wide ranges of B/Be ob-
served in arc volcanic rocks (5 to 200;
Morris etal, 1990).

In this report, B and Be concentration
data are presented for metamorphosed sedi-
mentary and mafic rocks and veins and
pegmatites from the Catalina Schist, ex-
posed on Santa Catalina Island in southern
California. Mineral reservoirs for these
elements during devolatilization of
metasedimentary and metamafic rocks are
considered, and the extent to which B and
Be were mobilized during progressive vola-
tile loss and migmatization is discussed.

Finally, the implications of these data for
trace element fractionation among miner-
als, H20-rich fluids, and silicate liquids
deep in subduction zones are discussed, as
well as the relevance of these observations
for interpretation of B-Be concentrations
in arc volcanic rocks.

The Catalina Schist consists of three
major metamorphic-tectonic units juxta-
posed along low-angle faults (Piatt, 1975;
Sorensen and Barton, 1987; Bebout and
Barton, 1989). These units contain similar
lithologies (metamorphosed sandstones,
shales, and cherts; metabasaltic and
metagabbroic rocks) and range in grade
from lawsonite-albite to amphibolite. The
Catalina rocks are well suited for examina-
tion of the effects of metamorphism on
trace element and stable isotope composi-
tion over a wide range of metamorphic
conditions (350°-750°C, 5-11 kbar). Evi-

ocmm m

Lawsonite-Albite
Blueschist

• o

Greenschist/Epidote
Amphibolite

(KCD Amphibolite

Seafloor Sediments
50 - 200 ppm B

80 120

Boron (ppm)

160

200

Fig. 9. Boron content in whole-rock metasedimentary samples from the Catalina Schist. The large
range in B concentration of the lawsonite-albite and blueschist grade rocks probably reflects
variability in sedimentary protoliths (see B data for seafloor sediments in Moran et al., 1991).
With increasing metamorphism, B content is decreased and becomes more uniform.

GEOPHYSICAL LABORATORY

dence for fluid transport and associated
mass transfer during metamorphism includes
the occurrence of veins, reaction zones be-
tween disparate lithologies, changes in bulk
chemical composition, and changes in iso-
topic composition (Bebout and Barton,
1989). Stable isotope and petrologic stud-
ies of the Catalina Schist have yielded evi-
dence for progressive devolatilization and
km-scale transport of H20-rich C-O-H-S-
N fluid during metamorphism (Bebout and
Barton, 1989; Bebout, 1991). Intheamphi-
bolite unit, pegmatites represent high-P mass
transfer via silicate liquids derived through
vapor-saturated partial melting of sedimen-
tary and mafic rocks (Sorensen and Barton,
1987; Bebout and Barton, 1989).

Boron and Beryllium Concentration Data

Boron and Be concentration data for 96
metasedimentary, metamafic, and meta-ul-
tramafic rocks, mineral separates, veins and
pegmatites were obtained by inductively
coupled plasma (ICP) analytical techniques;
samples were fused with Na2C03 flux. All
chemistry was done in the laboratory of J.
D. Morris and F. Tera at the Department of
Terrestrial Magnetism.

Boron concentrations and B/Be of
metasedimentary rocks decrease progres-
sively with increasing metamorphic grade
(Figs. 9, 10); the decrease in B content is
consistent with results for other metamor-
phic suites (e.g., Shaw etal., 1988; Nabelek
et al., 1990; Moran et ai, 1991). Lowest-
grade lawsonite-albite rocks (inferred meta-
morphic conditions of 350°-450°C, 5-8 kbar)
range in B content from 12 to 181 ppm B

with a mean of 73 ppm, whereas high-grade
amphibolite equivalents (inferred metamor-
phic conditions of 650°-750°C, 8-11 kbar)
range from 5.4 to 19 ppm B with a mean of
12.2 ppm. Beryllium contents in metasedi-
mentary rocks of all grades range from 0.3
to 1 .2 ppm. Thus, the decrease in B/Be from
a mean of -72 for the lawsonite-albite grade
metasedimentary rocks to a mean of -27 for
amphibolite grade equivalents is attributable
to loss of B (Fig. 10). Metamafic rocks
contain from 3 to 20 ppm B (12 samples)
and from 0.21 to 1.31 ppm Be, with the
lower values for both elements occurring in
the highest grade rocks. Metamafic B/Be
shows no significant variation with increas-
ing metamorphic grade and ranges from 3 to
20 (Fig. 10).

As measures of B and Be mobility in
hydrous fluids and felsic silicate melts, B
and Be concentrations were obtained for
veins, which precipitated from the hydrous
fluids, and for pegmatites, which reflect
migmatization in the amphibolite unit. In
blueschist metasedimentary exposures,

1000F

100

Lawsonite-Albite
Blueschist —-*■"" *» ^*

Greenschist
& Epidote
Atiphibolite

10 100

B (ppm)

1000

Fig. 10. B/Be vs. B content of metasedimentary
rocks and metamafic rocks (shaded field). B and
B/Be of metasedimentary rocks decrease with
increasing grade. With the exception of one
glaucophanic greenschist sample (60 ppm B, 1.98
ppm Be, B/Be of 30), all of the measured
metamafic rocks have B/Be less than 20.

CARNEGIE INSTITUTION

i 1 —

n = 10; 32.4
±10.3

Na-amphibole

(n = 10; 33.4 ±

11.5)

Blueschist Metasediments

Blueschist Veins

OO Host-Rocks

O Pegmatites

O Metasedimentary
Metamafic

B/Be

100

Fig. 11. Comparisons of B/Be of veins and pegmatites with those of host-rocks from the
blueschistand amphibolite units.

nearly monomineralic sodic amphibole
veins have B and Be contents and B/Be
ranges similar to their host rocks (Fig. 11).
One albite + graphite vein contains 4.6 ppm
B and 0.13 ppm Be and also has B/Be
(35.4) similar to that of the surrounding
blueschist grade metasedimentary rocks
(32.4 ±10). Pegmatites in the amphibolite
unit show wide ranges in B content (6 to 26
ppm) and Be content (0.6 to 4. 1 ppm) but
have B/Be similar to or slightly lower than
their mafic and sedimentary hosts (Fig. 11).
Consideration of B and Be behavior
during devolatilization and migmatization
requires knowledge of B and Be mineral
residency. Table 7 contains data for whole-
rock and mineral separate samples from
several high-grade metasedimentary rocks
and pegmatites. These data demonstrate
that B is concentrated in white micas,
whereas Be appears to be somewhat more
evenly distributed. For all but sample 7-3-
23, which contains tourmaline, muscovite

contains more B than the whole-rock
sample; Be content of the muscovite is
comparable to that of the whole-rock
samples. Preferential enrichment of B in
micas is consistent with the ion microprobe
results of Domanik et al. ( 1 99 1 ) for samples
of the Catalina Schist. With the exception
of several occurrences in felsic pegmatites
(amphibolite unit) and in greenschist-grade
metacherts. tourmaline is not a significant
host for B in the Catalina samples.

Discussion

The data presented in this report are
consistent with removal of B and Be from
metamorphosed sedimentary and mafic
rocks by both H20-rich fluids and felsic
silicate liquids. The B-Be signatures of
these two "fluids" were apparently dra-
matically different. The H20-rich fluids

GEOPHYSICAL LABORATORY

Table 7. Boron and beryllium concentrations of mineral separates (in ppm).

Whole-Rock

Muscovite

Feldspar

Sample

* not determined

** tourmaline-bearing pegmatite from metasedimentary exposure

Metasedimentary Rocks

6-3-41'

(epidote amphibolite)

32.0

4.1 n.d.* n.d.

4.2

1.1

8-1-3
(amphibolite)

19.0

0.41 39.1 0.69
Pegmatites - Amphibolite Unit

n.d.

6-3-24
6-3-75
7-3-23**

37.0
25.7
970

1.3 48.0 0.92
4.1 60.6 4.0
2.1 113 3.0

n.d.
11.1
n.d.

n.d.
5.8
n.d.

liberated through progressive
devolatilization are inferred to have had
high B/Be; their removal resulted in de-
crease of the B/Be of the residual rocks. To
explain the similarity of the Na-amphibole
vein and blueschist metasedimentary rock
B and Be concentrations and B/Be, a model
is suggested where veins precipitated from
fluids which were previously equilibrated
with respect to B and Be with host rocks (or
with similar rocks upstream of current hosts).
Stable-isotope data are consistent with this
model; O, H, and C isotope compositions of
vein minerals are in many cases similar to
those of the same minerals in host rocks
(Bebout and Barton, 1989). If this model is
correct, the veins would have had mineral/
fluid partition coefficients for B and Be
similar to the bulk-rock/fluid partition coef-
ficients of the host rocks. If the pegmatites
are directly representative of silicate liquid
compositions produced during melting, the
silicate liquids had B/Be similar to those of

host rocks. The relatively low B/Be of the
pegmatites may thus reflect earlier removal
of B from the source rocks during
devolatilization (Figs. 10, 1 1 ; cf. Leeman et
al, 1991).

Release of B and other fluid-mobile trace
elements during devolatilization may be
imagined as a result of discontinuous reac-
tions involving breakdown of mineral hosts
(e.g., see Nabelek et at., 1990), or by a
process of continual partitioning from min-
eral hosts into fluids, or by a combination of
these processes. ThegoodfitoftheCatalina
N concentration and isotope data with a
Rayleigh distillation model (Bebout and
Fogel, 1 991 ; Bebout, this Report) may have
implications for the mechanisms of B loss
during progressive devolatilization. Boron
and N show correlated decreases in concen-
tration with progressive volatile loss in the
Catalina metasedimentary rocks (see
Bebout, this Report), and both appear to be
concentrated in white micas. During pro-

CARNEGIE INSTITUTION

gressive de volatilization of the Catalina
Schist, N isotopic composition evolved
through incremental loss of N2 equilibrated
with the rock N reservoir, which was largely
dominated by the micas. Decreases in B
concentration may also reflect progressive
boron partitioning from B-rich minerals
(primarily the micas) into H^O-rich fluids
derived largely by chlorite-breakdown re-
actions (see Bebout, 1991). Thus, the B
loss may not simply reflect the breakdown
of micas, which show no variation in com-
bined modal abundance with grade.

These results virtually eliminate a bulk
sediment/slab mixing process (presumably
mixing of a melt derived from heteroge-
neous slab sources) as an explanation of B-
Be systematics in high-B/Be volcanic arcs
and strengthen the arguments of Morris et
al. ( 1 990) for addition to arc magma sources
of a fractionated (high-B/Be), slab-derived,
hydrous-fluid component. The data pre-
sented here predict that, before subducted
mafic and sedimentary rocks reach depths
of the Wadati-Benioff zone below arcs (80
to 150 km), the B/Be of these rocks is likely
to be decreased to <40. As is also suggested
by Moran et al. (1991), the simple mixture
of sediment and slab with B/Be reduced by
devolatilization will not produce a mixing
component with sufficiently high B/Be to
explain the linear trends in the arc data (see
Fig. 2 in Morris et al, 1990).

Because fractionation of B and Be in the
mantle wedge during melting processes is
unlikely (Morris et al, 1990; Ryan and
Langmuir, 1991), the varying impact of
hydrous fluid or melt additions to arc source
regions may explain some of the B/Be
variability observed in arc volcanic rocks

(see B-Be data for arcs in Morris et al.,
1990; Leeman et al, 1990; Ryan and
Langmuir, 1991). Subduction zones with
cooler inferred thermal structures on aver-
age show higher B/Be (e.g., Aleutian and
New Britain arcs with B/Be of 5-40 and 40-
200, respectively), consistent with addi-
tion of high B/Be hydrous fluids. Subduc-
tion zones with hotter inferred thermal struc-
tures produce volcanic rocks with rela-
tively low B/Be (e.g., the Cascades and
Woodlark Basin with B/Be of 2-10 and 4-
15, respectively), perhaps consistent with
addition of a melted component from a slab
previously stripped of B by lower-tempera-
ture devolatilization. Morris et al. (1990)
and Ryan and Langmuir (1991) have docu-
mented decreases in B/Be across individual
arcs, in all cases varying from high front-
arc values to low B/Be indistinguishable
from MORB in back-arc regions. These
decreases could presumably represent the
diminishing effects of hydrous fluid addi-
tion (contributing high-B/Be signatures)
and/or the onset of melt dominated B-Be
transfer (contributing low-B/Be signatures).

Conclusions

These results demonstrate that B and Be
are significantly fractionated during meta-
morphic devolatilization; this process pro-
duces high-B/Be hydrous fluids and results
in the dramatic reduction of the B/Be of the
subducted rocks. In contrast, partial melt-
ing, which does not strongly fractionate B
and Be, may produce silicate melts with
low B/Be inherited from previously

GEOPHYSICAL LABORATORY

devolatilized source rocks. During
devolatilization, loss of B and N occurred
as the elements partitioned from B- and N-
rich phases (e.g., micas) into H20-rich
fluids produced primarily by chlorite-
breakdown reactions. The B and Be con-
centrations of metasedimentary and
metamafic rocks, veins, and pegmatites of
the Catalina Schist place constraints on
models that invoke addition of these ele-
ments to arc source regions by "fluids"
(H20-rich solutions or silicate liquids).
The varying impact of hydrous fluid or
melt additions to arc source regions could
in part explain the variability in B/Be of
arc volcanic rocks, including cross-arc
variability in some individual arcs. Boron
and Be may serve as analogues for other
incompatible elements (e.g., Cs and Zr,
respectively) which appear to behave simi-
larly in hydrous fluid-solid and melt-solid
systems.

References

Bebout, G. E., Field-based evidence for
devolatilization in subduction zones: impli-
cations for arc magmatism, Science, 251,413-
416, 1991.

Bebout, G. E., and M. D. Barton, Fluid flow and
metasomatism in a subduction zone hydro ther-
mal system: Catalina Schist terrane, Califor-
nia, Geology, 17, 876-980, 1989.

Bebout, G. E., and M. L. Fogel, Nitrogen-isotope
compositions of metasedimentary rocks in
the Catalina Schist: implications for meta-
morphic devolatilization history, Geochim.
Cosmochim. Acta, in press, 1991.

Domanik, K., R. L. Hervig, and S. M. Peacock,
Beryllium and boron contents of subduction
zone minerals : an ion microprobe study, EOS,
72,293-294, 1991.

Leeman, W. P., D. R. Smith, W. Hildreth, Z.
Palacz, and N. Rogers, Compositional diver-

sity of Late Cenozoic basalts in a transect
across the southern Washington Cascades:
Implications for subduction zone magmatism,
Jour. Geophys.Res., 95, 19561-19582, 1990.

Leeman, W. P., V. B. Sisson, and M. R. Reid,
Boron geochemistry of the lower crust: Evi-
dence from granulite terranes and deep crustal
xenoliths, Geochim. Cosmochim. Acta, in press,
1991.

Moran, A. E., V. B. Sisson, and W. P. Leeman,
Boron in progressively metamorphosed oce-
anic crust and sediments: implications for com-
positional variations in subducted oceanic slabs,
Earth Planet. Sci. Lett., in press, 1991.

Morris, J. D., W. P. Leeman, and F. Tera, The
subducted component in island arc lavas: Con-
straints from Be isotopes and B-Be systemat-
ics, Nature, 344, 31-36, 1990.

Nabelek, P. I., J. R. Denison, and M. D. Glascock,
Behavior of boron during contact metamor-
phism of calc-silicate rocks at Notch Peak,
Utah, Amer. Mineral, 75, 874-880, 1990.

Piatt, J. P., Metamorphic and deformational pro-
cesses in the Franciscan Complex, California:
Some insights from the Catalina Schist terrain,
Geol. Soc. Amer. Bull, 86, 1337-1347, 1975.

Ryan, J. G., and C. H. Langmuir, Beryllium sys-
tematics in young volcanic rocks: Implications
for 10Be, Geochim. Cosmochim. Acta, 52, 237-
244, 1988.

Ryan, J. G., and C. H. Langmuir, The systematics
of boron abundances in young volcanic rocks,
Geochim. Cosmochim. Acta, in press, 1991.

Shaw, D. M., M. G. Truscott, E. A. Gray, and T. A.
Middleton, Boron and lithium in high-grade
rocks and minerals from the Wawa-
Kapuskasing region, Ontario, Can. Jour. Ear.
Sci.,25, 1485-1502, 1988.

Sorensen, S. S., and M. D. Barton, Metasomatism
and partial melting in a subduction complex:
Catalina Schist, southern California, Geology,
15, 115-118, 1987.

Tera, R, L. Brown, J. Morris, I. S. Sacks, J. Klein,
and R. Middleton, Sediment incorporation in
island arc magmas: Inferences from 10Be,
Geochim. Cosmochim. Acta, 50, 535-550, 1986.

CARNEGIE INSTITUTION

Laser Fluorination of Sulfide Minerals
with F2 Gas

D. Rumble, J. M. Palin, and T. C.
Hoering

Excitement pervades the discipline of
stable-isotope geochemistry at the pros-
pects for scientific advances raised by the
development of new, microanalytical and
in situ sampling techniques. Plans can now
be made, with reasonable chances of suc-
cess, to measure fundamental indicators of
mass transfer mechanisms in Earth's crust,
i.e. submillimeter-scale gradients in isoto-
pic compositions. Theoretical studies dem-
onstrate it is the curvatures, slopes, and
magnitudes of composition vs. distance
profiles that provide definitive evidence on
flux magnitude and direction and the rela-
tive importance of diffusive and infiltrative
mass transfer (Bickle and McKenzie, 1987;
Blattner and Lassey, 1989; Baumgartner
and Rumble, 1988; Bowman and Willett,
1991). Experience shows that advancing
the understanding of mass transfer in na-
ture depends on the ability to resolve differ-
ences in isotopic compositions over small
distances (Bickle and Baker, 1990a,b).
Isotopic microanalysis by laser-heating

minerals immersed in a reactive atmosphere
has been successfully demonstrated for sili-
cate, oxide, and sulfide minerals. Sharp
(1990) showed that laser heating of small
amounts of powdered quartz, olivine, po-
tassium feldspar, garnet, biotite, musco-
vite, diopside, and magnetite in a BrFs
atmosphere released O2 with 8^0 values
that are in good agreement with the results
of conventional BrF5 analyses. The in situ
analysis of spots on the surfaces of
uncrushed grains of quartz, plagioclase,
olivine, and magnetite has also been car-
ried out (Sharp, 1990; Schiffries and
Rumble, Annual Report 1989-1990, p. 37-
40; Elsenheimer et ai, 1990; Conrad and
Chamberlain, 1991). In situ analysis for
<534S in the minerals pyrite, pyrrhotite,
sphalerite, galena, and chalcopyrite has
been accomplished by laser heating and
combustion in an O2 atmosphere and
(Kelley and Fallick, 1990; Crowe et al.,
1990; Crowe and Shanks, 1991). The di-
ameter of analysis spots achieved in these
studies is 100-300 Jim, with a precision of
± 0.2%o.

We are currently testing the efficacy of
F2 gas in laser fluorination to analyze sul-
fide, silicate, and oxide minerals for ^4S
and ^l 80. It is desirable to develop alterna-

Table 8. Precision of laser fluorination using F2 gas in analyzing mineral powders

Sample

#4S

533S

#6S

CDT +0.16 (±0.13) 17

NBS-123 +17.62 (±0.15) 18

E40 -26.72 (±0.3) 14

M-pyrite +3.17 (±0.3) 33

+0.03

(±0.06)

-4.1

(±6.1)

+8.98

(±0.13)

+34

(±11)

-13.82

(±0.16)

-33.7

(±4.5)

+1.58

(±0.17)

+24

(±74)

o~= Standard Deviation

CDT: Canyon Diablo Troilite. NBS-123: Sphalerite. E40: Synthetic Ag2S prepared from
pyrrhotite; three wildly errant values for <5^6S have been dropped. M-pyrite: Pyritehedrons, 1
cm in diameter.

GEOPHYSICAL LABORATORY

tive fluorination reagents in order to opti-
mize analysis conditions. Conventional
analyses of silicates for <5180 show that
choosing BrFs or F2 may lead to incom-
plete fluorination and inaccurate analytical
results for certain minerals (Taylor and
Epstein, 1962; Clayton and Mayeda, 1963).
In the laser analysis of sulfides, it may
prove that laser fluorination gives less frac-
tionated analytical results than laser com-
bustion with O2 (cf. Kelley and Fallick,
1990; Crowe et ai, 1990). We report,
below, verification of the feasibility of F2
gas for micro analysis of 5^4S, cf^S, and
<536S in sulfide minerals.

Feasibility of Sulfide Laser Fluorination
with F2 Gas

The feasibility of using F2 gas as fluori-
nating agent to produce SF6 for the analysis
of 533s, #4S, and ^36s during laser heat-
ing of sulfide minerals has been established
by repeated analysis of aliquots of pow-
dered troilite, pyrite, sphalerite, and syn-

v>
«

dSJ

_ I

1 1 1 1 1 1

I I I | I I I I | I I I

*
* —

+ NBS-123

X CDT

•

—

O E40-SO2

—

O E40-BrF5

♦

—

.«*

-10

—

•

*
*

-20

—

™

— *

-*r*\ 1 1 1 i 1 1 1 1 I 1 1 1 1 I 1 1 1

Illl1

-30

-20

-10 0

534S, accepted

Fig. 12. Plot of accepted vs. measured values of
#4SforNBS-123 sphalerite, CDT-Canyon Diablo
Troilite, and E40 - synthetic Ag2S prepared from
pyrrhotite and analysed by conventional combus-
tion (SO2) or conventional fluorination(BrF5).

thetic Ag2S. The precision of analysis for
troilite and sphalerite is ± 0.06-0. \5%c for
S&S and #4S and± 6-1 Wcfor &£>S (Table
8). The precision for pyrite and synthetic
Ag2S is ± 0. 1 6-0 3%o for S^S and #4S and
± 27-74%o for ^S (Table 8). The preci-
sion for 534S is comparable to that obtained
by laser oxidation/combustion of sulfide
minerals in an O2 atmosphere (Crowe et
al., 1990). The scatter in the results for
pyrite may be explained, at least in part, by
the fact that the mineral was used for prac-

Table 9. Accuracy of laser fluorination using F2 gas. Mineral powders.

&*S

Sample

(measured)

(accepted)

(measured- accepted)

CDT

NBS-123

E40

E40*

+0.16
+17.62
-26.72
-26.72

0.0
+17.09
-27.12
-26.88

+0.16
+0.53
+0.40
+0.16

CDT: Canyon Diablo Troilite; standard #4S value defined asO.O. NBS-123: Sphalerite:
Value of +17.09 reported as average of results of intercomparsion of 1 1 laboratories.
(International Atomic Energy Agency, 1986). E40: Synthetic Ag2S prepared from
pyrrhotite. Accepted value analyzed by conventional combustion (Oliver et al., Annual
Report 1989-1990, p. 30-33). E40': Synthetic Ag2S prepared from pyrrhotite. Accepted
value analyzed by conventional BrFs fluorination (Oliver et al., Annual Report 1989-
1990, p. 30-33).

CARNEGIE INSTITUTION

2.0

_ I I I I | I I I I J...!...!...!...!...|...!...j...j....i...i...i....i....i.% i -

1.5 -

£ 1.0 ^

W « IT

S 0.5

-0.5

+ NBS-123

X CDT

O E40-SO2

O E40-BrFs (shift +2 along X-axis)

o.o E- Q

- 1 i i i i i i i i i i i i i i i i i i i i i i i-

-30

-20

-10 0

634S, accepted

Fig. 13. Plot of accepted values of $4S vs. the
difference between measured and accepted values
of^S.

tice in learning how to use F2 gas. The
scatter in results for synthetic Ag2S has not
yet been fully explained but may be related
to the relatively smaller amount of avail-
able sulfur as constrained by stoichiom-
etry. The accuracy of the results was evalu-
ated by comparing 5^4S of SF6 derived
from F2 laser fluorination of CDT, NBS-
123, and E 40 to "accepted" #4S values
obtained by conventional combustion with
O2 and fluorination with BrF5. The results
show deviations of from +0.16 to +0.53%c
in <P4S for F2 laser fluorination samples
relative to accepted values (Figs. 12, 13,
Table 9).

The experimental apparatus used for
F2 laser fluorination resembles, in general
outline, that used by Sharp (1990) for BrF5
fluorination of silicate minerals. There has
been added, however, a fluorine gas gen-
eration and delivery system, a heated KBr
trap to dispose of excess F2, and an Inconel
capacitance manometer to measure yields
and to monitor the progress of reactions
and the course of cryogenic transfers of
condensable gases. Fluorine is generated
by heating the compound K2NiF6 • KF in
an evacuated nickel reservoir to tempera-

tures of 290°-320°C (Asprey, 1976).* El-
emental fluorine is incorporated at low
temperatures (~250°C) and released at
higher temperatures according to the re-
versible reaction (Asprey, 1976)

2(K2NiF6 • KF) = 2K3N1F6 + F2 (1)
solid solid gas

Use of the compound solves two problems
encountered in fluorine chemistry. (1) F2
gas can be obtained free from contamina-
tion by N2 or O2. We outgas the compound
at 100°C under vacuum prior to evolution
of F2. (2) The hazard of handling large
quantities of F2 gas is eliminated. We
generate only the small amounts of F2
needed for laser fluorination. Excess F2
can be resorbed by the reversal of reaction
(1) at lower temperature.

The experimental procedure consists of
generating F2 gas, expanding it into an
evacuated sample chamber loaded with
samples, and aiming and firing the laser.
The reaction chamber is open to a liquid
nitrogen cold trap to remove SF6 cryogeni-
cally from the reaction site as soon as it is
formed. After laser heating, excess F2 is
removed by reaction with KBr. The prod-
uct SF6 is held back in a cold trap to avoid
mixing it with the bromine formed when F2
reacts with KBr. The apparatus is not fully
optimized for laser fluorination of sulfides
because at this step of the procedure prod-
uct SF6 must be removed in a Ni-metal
container and carried to another vacuum

1 We are indebted to J. O'Neil for alerting us to the
existence of the K2NiF6 • KF compound, to G. P.
Landis for furnishing unpublished information on
its use, and to Ozark- Mahoning, Inc. for manufac-
turing it.

GEOPHYSICAL LABORATORY

line for purification. The purification con-
sists of reacting SF6 with moist KOH to
eliminate traces of F2 and HF followed by
gas chromatography to remove trace con-
taminants that give isobaric interferences
with the ion beams of SF5+ in the mass
spectrometer (Hoering, Annual Report
1989-1990, p. 128-131).

Experimental parameters, such as pres-
sure of F2 gas and laser power, were varied
systematically to establish optimal operat-
ing conditions. Pressure of F2 gas in the
reaction chamber was varied from 55 to
175 torr. Fluorination proceeds readily at
lower pressures and is preferred to econo-
mize on consumption of the reagent. The
20-watt, CO2 laser was operated in both
continuous and pulsed mode. In the analy-
sis of powders, minimum power was used
in order to protect the fragile BaF2 win-
dows from damage. For most samples,
pulsed operation with a pulse spacing of 10
milliseconds and pulse width of 10 milli-
seconds at a laser power setting of 20% is
adequate to achieve complete fluorination.
An additional problem is encountered if the
initial laser shot is set at high power: the

10
5
0
-5

_i i i i | i i i i | i i i i | i i i i | i i i i_

.4?

j¥

<c*»

<*6

■ r^ • * *

-10 -

_ *

.15 Eg- i i i 1 i 1 1 i 1 1 1 1 i i i 1 i i 1 1 1 1-

►£• Sulfide Powders

-30 -20 -10

10 20

534S

Fig. 14. Plot of $3S vs. #4S showing measured
vs. theoretical mass fractionation.

powdered sample may be scattered around
the sample chamber. The problem may be
mitigated by starting the laser at low power
and low pulse width and increasing incre-
mentally until reaction is seen. It is diffi-
cult to avoid some scattering of powdered
samples. For this reason, yields of SF6 in
relation to weighed amounts of samples are
less than stoichiometric.

Analysis for four sulfur isotopes, 32s,
33S, 34S, and 36s, is made simultaneously
on aFinnigan-M AT '251 mass spectrometer
with custom quadruple collector. The mea-
surement of all the isotopes provides a
useful test for the quality of the analysis
(Fig. 14).

The theory of isotope fractionation in
chemical processes predicts that the ratio
<534SA533s = 1.94 (±0.01) in samples such
as terrestrial ores, where anomalous nucleo-
synthetic effects are absent (Hulston and
Thode, 1965). We obtained a value of
1.950 (± 0.003) for the ratio (Fig. 14).

The results reported above validate the
use of F2 gas in laser fluorination of sul-
fides. The laser fluorination of sulfides is
far quicker than conventional fluorination:
reactions are completed in seconds or min-
utes rather than overnight. In comparison
to laser combustion of sulfide minerals in
O2 (Crowe et al. 1 990), laser fluorination is
slower, because of the preparative gas chro-
matography that is required to eliminate
isobaric interferences in the mass spec-
trometer. The advantages of the method
are clear, however. One obtains precise
data on three of the four isotopes 32S, 33S,
and 34S. The additional uncertainties in-
troduced by oxygen isotope corrections in

CARNEGIE INSTITUTION

mass spectrometry of SO2 are eliminated,
for fluorine has only one stable isotope.
Because SF6 is chemically inert, does not
sorb onto the interior surfaces of vacuum
lines, and is not sensitive to moisture, it is
potentially a better working gas than SO2
for mass spectrometry of very small samples
(Rees, 1978).

References

Asprey, L. B., The preparation of very pure fluo-
rine gas, J. Fluorine Chem., 7, 359-361, 1976.

Baumgartner, L. P., and D. Rumble, Transport of
stable isotopes. I. Development of a kinetic
continuum theory for stable isotope transport,
Contrib. Mineral. Petrol, 98, 417-430, 1988.

Bickel, M., and D. McKenzie, The transport of
heat and matter by fluids during metamor-
phism, Contrib. Mineral. Petrol., 95, 384-392,
1987.

Bickel, J., and J. Baker, Advective-diffusive trans-
port of isotopic fronts: an example from Naxos,
Greece, Earth Planet. Sci. Lett., 97, 78-93,
1990a.

Migration of reaction and isotopic

fronts in infiltration zones: Assessments of fluid
flux in metamorphic terrains, Earth Planet. Sci.
Lett., 98, 1-13, 1990b.

Blattner, P., and K. R. Lassey, Stable-isotope
exchange fronts, Damkohler numbers, and fluid
torockratios, Chem. Geoi, 78, 381-392, 1989.

Bowman, J. R, and S. D. Willett, Spatial patterns
of oxygen isotope exchange during one dimen-
sional fluid infiltration, Geophys. Res. Lett., 18,
971-974, 1991.

Clayton, R.N., and T. K. Mayeda, The use of BrF5
in the extraction of O2 from oxides and silicates
for isotopic analysis, Geochim. Cosmochim.
Acta, 27, 43-52, 1963.

Conrad, M. E., and C. P. Chamberlain, Laser-
based analyses of small-scale variations in the
oxygen isotope ratios of hydrothermal quartz,
£05,72,292,1991.

Crowe, D. E., J. W. Valley, and K. L. Baker,
Micro-analysis of sulfur-isotope ratios and zo-
nation by laser microprobe, Geochim.
Cosmochim. Acta, 54, 2075-2092, 1990.

Crowe, D. E., and W. C. Shanks, Laser micro-
probe $4S study of coexisting sulfide pairs:
seeing through metamorphism, EOS, 72, 1991.

Elsenheimer, D., J. W. Valley, and K. Baker, In-
situ laser microprobe determinations of (5180,

GSAAbstractsw. Programs, 22, 160-161, 1990.

Hulston, J. R., andH. G., Thode, Variations in the
33S, 34S, and 36S contents of meteorite and their
relation to chemical and nuclear effects, /.
Geophys. Res., 70, 3475-3484, 1965.

Kelley, S. P.. and A. E. Fallick, High precision
spatially resolved analysis of £*4S in sulfides
using a laser extraction technique, Geochim.
Cosmochim. Acta, 54, 883-888, 1990.

Rees, C. E., Sulfur isotope measurements using
SO2 and SF6, Geochim. Cosmochim. Acta, 42,
383-389, 1978.

Rumble, D., T. C. Hoering, and J. M. Palin, Mi-
croanalysis for $AS in sulfide minerals with
laser fluorination, EOS, 72, 292, 1991.

Sharp, Z. D., A laser-based microanalytical method
for the in situ determination of oxygen isotope
ratios of silicates and oxides, Geochim.
Cosmochim. Acta, 545, 1353-1357, 1990.

Taylor, H.P., and S. Epstein, Relationships be-
tween 180/160 ratios in coexisting minerals of
igneous and metamorphic rocks, Geol. Soc.
Amer. Bull., 73, 461-480, 1962.

Stable Isotope and Trace Element Indi-
cators of Devolatilization History in
Metashales and Metasandstones

Gray E. Bebout

Devolatilization of carbonate-poor
metashales and metasandstones has the po-
tential to release large amounts of H2O-
rich C-O-H-S-N fluid (e.g., Walther and
Orville, 1982). However, few studies have
directly examined this process, in contrast
with the many studies that have dealt with
devolatilization/inf iltration in meta-carbon-
ate systems (e.g., Valley, 1986). Because
metashale and metasandstone make up a
significant fraction of the continental crust,
their devolatilization may be extremely
important for crustal chemical, thermal,
and rheological evolution. The
metasedimentary rocks of the Catalina
Schist (California) are well suited for a

GEOPHYSICAL LABORATORY

study of devolatilization; rocks of similar
bulk composition have been metamor-
phosed under a wide range of metamor-
phic conditions (350°-750°C, 5-11 kbar;
Piatt, 1975; Sorensen and Barton, 1987;
Bebout and Barton, 1989). Trace element
and stable isotope compositions, mineral
modes, and mineral compositions show
distinct covariance with increasing meta-
morphic grade. In this study, the integra-
tion of petrologic and geochemical data
results in a distillation model for progres-
sive devolatilization. This model neces-
sarily implies specific devolatilization
mechanisms which have consequences for
models of crustal chemical evolution, fluid
transport dynamics, and metamorphic re-
action kinetics.

The three major metamorphic units of
the Catalina Schist (lawsonite-albite/
blueschist, glaucophanic greenschist/epi-
dote-amphibolite, and amphibolite) con-
tain sedimentary, mafic, and ultramafic
rocks underplated and metamorphosed
during early Cretaceous subduction (Piatt,
1975; Bebout, 1991a). Metasedimentary
rocks from the three units show a similar
range in lithology from metapelites to meta-
graywackes; average grain size increases
with increasing grade (several \im to sev-
eral mm). Trends in H2O content with
increasing metamorphic grade and pro-
grade reaction histories inferred from min-
eral modes indicate that devolatilization of
the metasedimentary rocks of the Catalina
Schist principally involved chlorite break-
down reactions over the approximate tem-
perature interval 400°-600°C (Bebout,
1991a). H2O content, determined by H-
isotope extraction techniques, decreases

from 2-5.5 wt % in lowest-grade rocks to
1.5-2.5 wt % in highest-grade, amphibolite-
facies rocks (Bebout, 1991a). Chlorite
breakdown reactions resulted in the produc-
tion of muscovite-, biotite-, garnet-, and
kyanite-bearing mineral assemblages
through reactions of the following general
types:

2Phengite + Chlorite = Muscovite

+ Biotite + Quartz + 4H2O (1)

2Chlorite + 4Quartz

= 3Garnet + 8H2O (2)

3 Chlorite + 7 Muscovite + Quartz

= Al2Si05(Kyanite) + 7Biotite

+ 12H20 (3)

Chlorite + Muscovite = Biotite +
Al2Si05(Kyanite) + Quartz + 8H2O (4)

In addition to the decrease in H2O con-
tent, devolatilization resulted in preferential
loss of some trace elements and in shifts in
the C and N isotope compositions of the
rocks. The metasedimentary rocks show
trends of decreasing N concentration and
increasing whole-rock 8 5N with increas-
ing metamorphic grade (Bebout and Fogel,
Annual Report 1989-1990, p. 19-26; Fig.
15). The £13C of carbonaceous matter in
these rocks increases from values of from
-26 to -24 %o in the lowest-grade rocks to
values of from -21 to -19 %o in the highest-
grade, amphibolite-facies rocks (Fig. 15).
Concentrations of carbonaceous matter do
not vary systematically with increasing grade
and average -0.6 wt % for all grades. De-
spite metamorphism at temperatures ex-

CARNEGIE INSTITUTION

■18

-20 -

§ -22

-24

-26

-28

Lawsonite- •
Albite &
Blueschist

Greenschist &

Epidote

Amphibolite

j i_

515N

4
Air

Fig. 15. Whole-rock 515N vs. <513C of carbon-
aceous matter for metasedimentary rocks of the
Catalina Schist. Note parallel trends toward higher
isotope values.

ceeding 350°C, C/N of the lowest-grade
metasedimentary rocks (5-20; mean -13)
is similar to C/N of many unmetamorphosed

sedimentary rocks, including those in trench
and off-trench environments (cf. Patience
et al., 1990). Higher-grade rocks have
higher C/N that is attributable to preferen-
tial N loss (for 6 greenschist and epidote
amphibolite samples, range is from 5 to
1 25; for four amphibolite samples, range is
from 28 to 237). Boron concentration in
the same rocks decreases with increasing
metamorphic grade from an average of -73
ppm in lawsonite-albite grade rocks to an
average of -12 ppm in amphibolite -grade
rocks (Fig. 16; Bebout et a/., this Report).
Bebout and Fogel (Annual Report 1989-
1990, p. 19-26) interpreted the N concen-
tration and isotope data to be results of
Rayleigh distillation behavior during pro-
gressive devolatilization (see equation [4]
in Valley, 1986). A remaining difficulty in
the Rayleigh distillation calculations stems
from variability in composition within an
individual grade. This variability, which in

1000

800

^ 600

400 -

200

Lawsonite-Albite &
Blueschist (• )

Greenschist & Epidote
Amphibolite (A )

Amphibolite (o )

100

Boron (ppm)

200

Fig. 16. N and B concentrations of metasedimentary rocks of the Catalina Schist.
See discussion of the significance of B loss for subduction zone recycling in
Bebout et al. (this Report).

GEOPHYSICAL LABORATORY

CL
Q.

C
CD
D)
O

800

600

400

200

Mean F Calculated

from Sample Pairs

0.26 + 0.07

(n = 12)

Lawsonite-Albite &
Blueschist

F 0.4 -

K>0 (weight %)

Fig. 17. Relationship of N and K2O concentra-
tions in metasedimentary rocks of the Catalina
Schist. Tie lines connect samples with similar
K2O content and demonstrate that high-grade
samples contain 0.26 ± 0.07 of the N in low-grade
samples with similar K content; calculated F is
shown for several of the tie lines. The mean 8 N
shift between low- and high-grade samples in
these same pairs is 1.85%c (standard deviation of

1.0%©).

part represents protolith variability, compli-
cates comparisons of trace element compo-
sition and volatile content of low- and high-
grade metamorphic equivalents. The
protolith variability problem can be allevi-
ated by normalization of the N data to K2O
data for the same rocks. Because N is
preferentially partitioned into micas (Honma
and Itihara, 1981), and because the micas
constitute the only significant K reservoir in
the rocks, the concentrations of N correlate
well with K2O content of the rocks (Fig.
17). By comparing only samples with simi-
lar K2O content (i.e., with similar modal
abundance of micas), fluid-rock isotope frac-
tionation factors can be more tightly con-
strained, as shown in Fig. 18.

Fluid - Rock

Fig. 18. Demonstration of the interdependencies
of 8 N shift due to Rayleigh distillation (labelled
curves), fluid-rock N-isotope fractionation
(A15Nfiuid-rqck; related to alpha, the fractionation
factor, by 8 ^Ifiuid - 8 Nrock = 103 In alpha), and
inferred N loss (F indicates the fraction of the
original N remaining in the rock.) Indicated are
estimates of N loss % based on differences in
mean N content in the lawsonite-albite and am-
phibolite facies rocks (0.22; dashed line) and
based on normalizations by use of the K2O data
indicates an F or ~0.26±0.07; see solid horizontal
lines that indicate mean ± one standard devia-
tion). Curves are for the mean of the 8 N shifts
for the pairs with similar K2O content (± one
standard deviation on lower and higher 8 N
sides; 1.85 ± 1 .0 %p- solid curves), and the differ-
ence in the mean 8 N of the lawsonite-albite and
amphibolite samples (dashed curve labeled 2.4%o;
see Bebout and Fogel, Annual Report 1989-1990,
p. 19- 26 for data). Shaded region indicates range
of fractionations (-1.511.0 %6) compatible with
the Catalina data based on these data for N loss
and 8 N shift. Vertical arrows indicate frac-
tionations calculated by use of differences in
mean N and 8 N of the low- and high-grade
rocks (—1.6 %o) and by use of the K^O-normal-
ized data (—1.4 %o).

Values of A15Nfiuid-rock (<515Nfiuid -
<515Nrock) of about -1 .5+1 .0%o inferred in
Fig. 1 8 are similar to values of from -1 .0 to
-4.0%o suggested by Haendel et ai (1986)
based on isotope and concentration shifts
in metamorphic suites, but are smaller

CARNEGIE INSTITUTION

than ranges of from -3.0 to -4.0 %oand -6.0
to -11.5 %o obtained by Kreulen et al
(1982) and Bottrell et al (1988), respec-
tively, for fluid inclusion-mineral/rock pairs
in low to medium grade metasedimentary
rocks. The range of fractionation factors
indicated in this study is similar to the
range calculated by Hanschmann (1981)
for N2-NH4"1" exchange equilibrium at 400°-
600°C (A15Nfiuid-rock of from -2.25 to -2.9
%o for the temperature range of 400°-
600°C). Speciation of N as N2 in the
Catalina fluids is further suggested by the
presence of small amounts of N2 in fluid
inclusions from metasomatized eclogitic
blocks in blueschist melange (<1 mol %; T.
C. Hoering, personal communication, 1991;
determined by quadrupole mass spectrom-
etry). Furthermore, calculations of fluid C-
O-H-N equilibria indicate that N2 is the
dominant N species under most crustal
metamorphic conditions (Duit etal, 1986;
Ferry and Baumgartner, 1987; Bottrell et
al., 1988).

Carbon and O isotope data for the same
rocks are compatible with the model of
Rayleigh distillation derived from the N
data. Progressive fractionation of 12C from
carbonaceous matter into CH4 in fluids by
a Rayleigh distillation process could pre-
sumably explain the observed shift in <513C
(Fig. 15; seeBebout, 1989). Because of the
relatively large A13C of CH4-graphite ex-
change at these temperatures (-3 to -7 %o\
see Bottinga, 1969), such a shift in £13C
could be achieved without a large decrease
in whole -rock C concentration. Heating/
freezing experiments (Sorensen and Barton,
1987) and quadrupole mass spectrometry
indicate that CH4 is the dominant C species

(<1-10 mol %) in fluid inclusions from
metasomatized mafic blocks in blueschist
and amphibolite unit melange. The unifor-
mity of <5180 of the metasedimentary rocks
at all grades (quartz <5180 of approximately
+16 to +19 %o; Bebout, 1991a) indicates
that the metasedimentary rocks did not
pervasively reequilibrate with the H2O-
rich fluids that produced veins and infil-
trated more permeable portions of the
Catalina Schist (Bebout, 1991a, b). This
inferred relative impermeability of the
metasedimentary rocks supports the as-
sumption of closed- system behavior im-
plicit in the Rayleigh distillation calcula-
tions.

The Rayleigh calculations assume that
isotopic equilibrium is maintained between
the fluid phase and the remaining mineral
phases and among mineral phases during
incremental loss of fluid. In the Catalina
metasedimentary rocks, the mechanisms
affording this continual reequilibration pre-
sumably involved diffusive exchange and
dissolution/reprecipitation during grain
coarsening. The increase in white mica
grain size, a trend in white mica chemistry
{decrease in celadonite substitution, [(Mg,
Fe2+) + Si = 2A1]; Sorensen, 1986; Bebout,
1989}, and increases in the grain size and
degree of crystallinity of carbonaceous mat-
ter (Bebout 1 989) presumably reflect these
processes. For other rocks, Duit et al
(1986) reported coupled decreases in mica
and whole-rock N concentration with in-
creasing metamorphic grade. Such de-
creases are consistent with progressive par-
titioning of N into fluids equilibrated with
the remaining micas, as opposed to selec-
tive release of N due to mica breakdown

GEOPHYSICAL LABORATORY

(i.e., with the remaining mica retaining its
original higher N content) . Similarly, in an
ion microprobe study of the rocks ana-
lyzed in this study, Domanik et al. (1991)
report a general decrease in the B concen-
tration of white micas with increasing meta-
morphic grade.

The proposed model requires that flu-
ids released by devolatilization escape
without appreciable infiltration by exter-
nally-derived fluids significantly out of
0-, C-, or N-isotopic equilibrium with the
rocks. Such a scenario might arise as
locally-derived fluids migrate within lay-
ers of similar composition and at similar
P-Tconditions toward fractures along pore
fluid pressure gradients (cf. Etheridge et
al., 1984). Veins, which represent the
larger fractures, are abundant in the Catalina
Schist, particularly at lower grades (Bebout
and Barton, 1989). The fluids in this
scenario would always be equilibrated with
the rocks along their flow paths and would
not leave a significant imprint on trace
element and stable isotope compositions
in the host rock. However, flow across
layering (e.g., of mixed sandstone and
shale sequences) would possibly homog-
enize elemental and isotopic variations
that resulted from local devolatilization
history or protolith variability, depending
on the magnitude of the variations, fluid-
to-rock ratios, and fluid composition. The
degree of homogenization would depend
in part on fracture density, which would
control the scale of the intergranular flow,
and the relative permeabilities of the
interlayered lithologies, which would dic-
tate whether or not intergranular flow oc-
curred across layering. To address further

these issues, ongoing research involves
analyses of trace element concentrations
and stable isotope composition in finely
interlayered, veined sequences of lawsonite-
albite and blueschist grade metasandstone
and metashale. Preliminary data indicate
the preservation of -1.0 %o gradients in
8 5N within one finely interlayered (cm-
scale) lawsonite-albite grade meta-sedimen-
tary exposure.

For carbonate-poor metamorphosed
sandstones and pelitic rocks, N and C (and
perhaps S) isotope systematics appear to be
more useful measures of the extent of
devolatilization than either the H or O sys-
tems, owing to a more favorable mass-
balance between the fluids and rocks. Be-
cause of the large reservoir of H in high-
X(H20) fluids and the low H/O of the rocks,
trends in H isotope compositions due to
fluid loss are probably modified more easily
by interaction with the small amounts of
fluid derived within local fluid generation/
extraction systems. The Catalina meta-
sedimentary rocks show no obvious trend in
SD with increasing metamorphic grade and
range from -80 to -50 %oat all grades (Bebout,
1989). Because of the large O reservoir in
the rocks, Rayleigh distillation results in
only minor change in whole-rock 5 °0 with
progressive devolatilization; this change is
probably unresolvable given the marked
protolith variability in <5180. Similar argu-
ments were made by Oliver et al. (Annual
Report 1989-1990, p. 30-33) for S-isotope
systematics in sulfidic schists from the
Waterville-Augusta area, Maine. Because
of the low S content in the metamorphic
fluids (2-5 mol %; Ferry, 1981), greenschist-
grade metamorphic rocks retained S-iso-

CARNEGIE INSTITUTION

tope signatures inherited from diagenetic
and lower-grade metamorphic de-
volatilization histories. These arguments
are also analogous to those made for O and
C isotope systematics during decarbon-
ation of impure carbonates infiltrated by
H20-rich, C-poor fluids. Because of the
low C content of the infiltrating fluids, C-
isotope composition of the rocks is re-
garded as a better measure of the extent of
decarbonation than the O-isotope system-
atics, which are commonly controlled by
infiltration (Rumble, 1982; VaUey, 1986).
A distillation mechanism like that pro-
posed for the Catalina N data may dictate
the mobilities of other fluid-mobile trace
elements (e.g., B, Cs, U; see Heier, 1973;
Leeman et aL, 1990) residing in a range of
mineral reservoirs (i.e., not only in micas,
which are the predominant N and B reser-
voir). The efficiency with which equilib-
rium is maintained between fluids and min-
erals during incremental loss may be dic-
tated by the relative intracrystalline diffu-
sion rates of the elements in their respective
mineral reservoirs or by dissolution/
reprecipitation rates of the host minerals
during fluid loss. Boron concentration,
like N concentration and isotope composi-
tion, may have been controlled by the bulk
fluid loss and fluid-rock exchange history
of the rocks. If so, B partitioned progres-
sively from B-rich minerals (primarily the
white micas) into H20-rich fluids derived
largely by chlorite breakdown. Thus, as
with N, the B loss need not be attributed to
breakdown of host minerals such as white
mica and biotite, which show no obvious
decrease in their combined modal abun-

dance with increasing metamorphic grade
[see equations ( 1 -4)] . Continuous exchange
reactions like those which stabilized micas
during progressive metamorphism may also
allow the continual reequilibration of fluid
and rock stable isotope and trace element
composition during progressive
devolatilization and/or infiltration by ex-
ternally-derived fluids.

References

Bebout, G. E., Geological and geochemical inves-
tigations of fluid flow and mass transfer during
subduction-zone metamorphism, Ph. D. Dis-
sertation, University of California, Los Ange-
les, 1989.

Bebout, G. E., Field-based evidence for
devolatilization in subduction zones: Implica-
tions for arc magmatism, Science, 251, 413-
416, 1991a.

Bebout, G. E., Geometry and mechanisms of fluid
flow at 15 to 45 kilometer depths in an early
Cretaceous accretionary complex, Geophys.
Res. Lett., 18, 923-926, 1991b.

Bebout, G. E., and M. D. Barton, Fluid flow and
metasomatism in a subduction zone hydrother-
mal system: Catalina Schist terrane, Califor-
nia, Geology, 17, 876-980, 1989.

Bottinga, Y., Calculated fractionation factors for
carbon and hydrogen isotope exchange in the
system calcite-C02-graphite-methane-hydro-
gen and water vapor. Geochim. Cosmochim.
Acta 33, 49-64, 1969.

Bottrell, S. H., L. P. Carr, and J. Dubessy, A
nitrogen-rich metamorphic fluid and coexist-
ing minerals in slates from North Wales, Min-
eral. Mag., 52, 451-457, 1988.

Domanik, K., R. L. Hervig, and S. M. Peacock,
Beryllium and boron contents of subduction
zone minerals: an ion microprobe study, EOS,
72,293-294,1991.

Duit, W., Jansen, J. B. H., Van Breemen, A., and
A. Bos, Ammonium micas in metamorphic
rocks as exemplified by Dome de L'Agout
(France). Amer. J. Sci., 286, 702-732, 1986.

Etheridge, M. A., Wall, V. J., Cox, S. R, and R. H.
Vernon, High fluid pressures during regional
metamorphism and deformation: Implications
for mass transport and deformation mech-
anisms,/ Geophys.Res.,89, 4344-4358, 1984.

GEOPHYSICAL LABORATORY

Ferry, J. M., Petrology of graphitic sulfide-rich
schists from south-central Maine: An example
of desulfidation during prograde regional meta-
morphism. Amer. Mineral., 66, 908-930, 198 1 .

Ferry, J. M., andL. Baumgartner, Thermodynamic
models of molecular fluids at the elevated pres-
sures and temperatures of crustal metamor-
phism, Rev. Mineral., 17, 323-365, 1987.

Haendel, D., K. Muhle, H.-M. Nitzsche, G. Stiehl,
and U. Wand, Isotopic variations of the fixed
nitrogen in metamorphic rocks, Geochim.
Cosmochim. Acta, 50, 749-758, 1986.

Hanschmann, G., Berechnung von Isotop Effekten
auf quantenchemischer Grundlage am Beispiel
stickstoffhaltiger Molekule, Zfl-Mitt., 41, 19-
39, 1981.

Heier, K. S., Geochemistry of granulite facies
rocks and problems of their origin, Phil. Trans.
R. Soc. Lond. A., 273, 429-442, 1973.

Honma, H., and Y. Itihara, Distribution of ammo-
nium in minerals of metamorphic and granitic
rocks, Geochim. Cosmochim. Acta, 45, 983-
988, 1981.

Kreulen, R., A. Van Breeman, and W. Duit, Nitro-
gen and carbon isotopes in metamorphic fluids
from the Dome de L'Agout, France, Proceed-
ings of the Fifth International Conference on
Geochronology, Cosmochronology, and Iso-
tope Geology, p. 191, 1982.

Leeman, W. P., A. E. Moran, and V. B. Sisson,
Compositional variations accompanying meta^
morphism of subducted oceanic lithosphere:
Implications for genesis of arc magmas and
mantle replenishment, Abstr., Seventh ICOG
Mtg., 58, 1990.

Patience, R. L., C J. Clayton, A. T. Kearsley, S. J.
Rowland, A. N. Bishop, A. W. G. Rees, K. B.
Bibby, and A. C. Hopper, An integrated bio-
chemical, geochemical, and sedimentological
study of organic diagenesis in sediments from
Leg 1 12, in Proceedings of the Ocean Drilling
Program, Scientific Results, 112, E. Suess et
al., eds., Ocean Drilling Program, College Sta-
tion, Texas, 135-153, 1990.

Piatt, J. P., Metamorphic and deformational pro-
cesses in the Franciscan Complex, California:
some insights from the Catalina Schist terrain,
Geol. Soc. Amer. Bull, 86, 1337-1347, 1975.

Rumble D., Stable isotope fractionation during
metamorphic volatilization. Rev. Mineral. 10,
327-353, 1982.

Sorensen, S. S., Petrologic and geochemical com-
parison of the blueschist and greenschist units
of the Catalina Schist terrane, southern Califor-
nia, Geol. Soc. Amer. Mem., 164, 59-75, 1986.

Sorensen, S. S., and M. D. Barton, Metasomatism
and partial melting in a subduction complex:
Catalina Schist, southern California, Geology,
15, 115-118, 1987.

Valley, J. W., Stable isotope geochemistry of meta-
morphic rocks, Rev. Mineral. 16, 445-489,
1986.

Walther, J. V., and P. M. Orville, Volatile produc-
tion and transport in regional metamorphism,
Contrib. Mineral. Petrol, 79, 252-257, 1982.

The fa Content of Normative ol
Felix Chayes

Replacement of conventional "wet-
way" analytical procedures by instrumented
physical techniques is generating a large
and rapidly expanding reservoir of rock
analyses in which Fe is not partitioned by
oxidation state. (Of the 14,722 analyses of
igneous rocks included in the current ver-
sion of the base IGBADAT, for instance,
almost a quarter lack Fe partition.) Com-
parison of analyses subject to this defi-
ciency with older data, indeed, plotting
them in many common petrographic varia-
tion diagrams, requires some external
nonanalytical assignment of Fe oxidation
state.

In the several procedures now in use
(for a thorough review see Middlemost,
1989), this adjustment consists of the ap-
plication of a simple formula to convert the
lone Fe or Fe-oxide value into two. With
few exceptions the conversion is made
without regard to the amounts of other
components reported in the bulk analysis.
And however important petrologic and
mineralogical factors may have been in the
development of these relations, they play
no direct role in their application. In prac-
tice, the preferred rule is simply applied by
rote to any analysis that is to be used in a
fashion requiring some estimate of the oxi-

CARNEGIE INSTITUTION

dation state of Fe in the specimen in ques-
tion.

In olivine1 bearing rocks this may lead
to ol unrealistically rich in fa. Fortunately,
loss of information about the oxidation
state of Fe in bulk analyses of igneous
rocks has been accompanied by a manifold
increase in the amount of information about
the chemical compositions of their con-
stituent minerals. It has been pointed out
(Chayes, Annual Report 1989-1990, p. 40-
42) that whenever such information is in
fact available for an analyzed specimen, it
would be a simple matter to adjust the Fe
oxidation ratio inversely, by using the fa
content of modal olivine, or some multiple
of it, as an upper limit on the normative
ratio fa/ ol. What does one do, however, if,
as is still true of many rocks and will always
be true of some, no pertinent mineralogical
information is available?

Rather than fall back on a single rule for
all rocks, one might then prefer an inverse
adjustment based on the average/a/0/ con-
tent found from more complete analyses of
rocks similar to that in question. This note
presents relevant summaries drawn from
IGBADAT. Of the 14,722 specimen de-
scriptions in the current version of that
base, 11,294 are accompanied by "com-
plete" bulk analyses, which sum to be-
tween 95 and 105% and include determina-
tions for both oxides of Fe. The cumulative
frequency distribution offal ol in the 4,589
of these that are 6>/-normative is shown by
the upper line in Fig. 19. The lower line
shows the same information for the 3,476

1 In this note, "olivine" denotes the minerall, ol and
fa the normative components computed from the

analysis

0.3 0.4 0.5 0.6

Lower Class Mark

Fig. 19. Cumulative frequency distributions of fa/
ol in analyses with summations in the range 95-
105 and containing analytical determinations for
both oxides of Fe. Data from IGB ADAT4 (4,589
analyses in all, 3,476 with H2O <2%, class width
= 0.05).

which also contain less than 2% of H2O.
Values of fa/ ol in excess of 0.5 comprise no
more than 2.5% of the entries in either data
set. In fact, in only 6% of all values
included in either summary is fa/ol greater
than 0.4, and in only 12% is it greater than
0.35.

The mean and standard error of this
statistic in each of a number or rock types
(really name groups) are shown in Table
10. Despite the limited overall range of fa/
ol and its rather broad within group disper-
sion, there appear to be marked differences
between group means; in randomly drawn
samples several of these differences would
be considered statistically significant. In
fact, with the exception of the tephrites and
gabbros the groups seem to fall into three
sets, with mean values of fa/ol in the
ranges, respectively, 0.28-0.31, 0.20-0.21
and 0.11-0.13. In Table 1 0 the members of
each set are listed in order of decreasing
size of the sample available for calculation
of the mean and its error.

As of this writing, calculations have
been completed only for complete groups;
from prior experience it is anticipated that

GEOPHYSICAL LABORATORY

Table 10. Average and standard error of fa/ ol, by rock type, data from IGBADAT4

Rock Type

Number of Analyses

falol

All

ol >0.5%

mean

std. error

Tholeiite

339

202

0.306

0.005

Dolerite

359

140

0.285

0.009

Diorite

358

0.310

0.013

Diabase

200

0.298

0.015

Andesite

794

0.281

0.021

Basalt

2,532

1,629

0.206

0.003

Gabbro

504

342

0.181

0.007

Basanite, hawaiite, mugearite

384

320

0.197

0.007

Trachyandesite, benmoreiite, tristainite

246

158

0.208

0.012

Phonolite

319

116

0.208

0.019

Mafic plutonics, ultramafic nodules

463

384

0.114

0.005

Ultramafic volcanic s & dikes

313

224

0.128

0.008

Tephrite

101

0.156

0.011

elimination of hydrated materials will shift
means upward slightly, and since in all
these groups high values of falol are either
very rare or lacking, this may well reduce
observed differences between means. It
should also materially reduce within group
dispersion, however, so that its effect on the
significance of intergroup differences re-
mains to be seen.

The underlying situation, alas, is prob-
ably not as simple as Table 10 may make it
seem. Even ignoring the evident unruli-
ness of the tephrites and gabbros, elaborate
and sometimes rather elliptical rationaliza-

tions about qualifications for membership
in some of the groups cannot be altogether
avoided in any organization of data based
on the actual usage of rock names rather
than on their a priori definitions. A more
thorough discussion of the problems this
raises is in preparation.

References

Middlemost, E.A.K., Iron oxidation, norms, and
the classification of igneous rocks, Chem.
GeoL, 77, 19-26, 1989.

GEOPHYSICAL LABORATORY

Igneous and Metamorphic Petrology
Experimental studies

Raman Spectra of High-Temperature

Silicate Melts: NA2O-S1O2, K2O-S1O2,

and L12O-S1O2 Binary Compositions

John D. Frantz and Bjorn O. My sen

The structure of silicate liquids, deter-
mined at high temperature, and relation-
ships between structure and properties are
centrally important to our understanding of
natural magmatic processes. Principally
from studies of quenched melts, for the
compositional range of most natural mag-
matic liquids, a simple equilibrium of the
form;

T205(2Q3) *=> T03(Q2) + TOiiQ4),1 (1)

describes the principal elements of the struc-
ture (e.g., Virgo et ai, 1980; Matson et al.,
1983; Stebbins, 1988), where T represents
tetrahedrally coordinated cations. From
spectroscopic data on quenched binary
metal oxide-silica melts (glasses), it ap-
pears that equation (1) shifts to the right
with increasing Z/r2 of the network-modi-
fying metal cation.

1 In this paper, stoichiometric units are used to
describe the structural units. T represents tetrahe-
drally coordinated cation(s) such as, for example,
Si4+ and Al-K In view of the frequent usage of Q-
notations in the NMR literature, the equivalent
notations are shown here for convenience. The
superscript in the Q-notation refers to the number
of bridging oxygens on the unit.

Whether this relationship holds true in
the molten state is not known. Studies
using NMR (Stebbins, 1988; Brandriss and
Stebbins, 1988) and Raman spectroscopy
(Seifertertf/., 1981; My sen, 1990; Cooney
and Sharma, 1990) indicate systematic
structural changes with temperature. There-
fore, even though the spectra of glasses and
melts qualitatively resemble one another,
at least for the limited compositions stud-
ied, structural data from glasses may not be
used to characterize relationships between
melt properties and melt structure. In-situ,
high-temperature spectra of melts are re-
quired. Measurements of the Raman spec-
tra of melts in their molten state to tempera-
tures of above 1600°C are now possible. In
the present study, the effect of temperature
on the spectra of the Na20-Si02, K2O-
Si02, and Li02-Si02 are investigated.

The integration of a micro-heating stage
with the focusing capability of the micro-
Raman system is fundamental to the suc-
cess of the proposed research. The heater/
thermocouple is fabricated from pieces of
0.8-mm Pt and Pt9oRhiO wire which are
welded and flattened to 500 jum thickness
at the join. A 1-mm diameter hole is drilled
through the Pt-PtooRh 10 junction (Fig. 20).
In the heating stage, the Pt-PtooRhio ther-
mocouple serves a dual function as both
thermocouple and heater (Ohashi and
Hadidiacos, Year Book 75, 828-834). The
heater responds to applied power in a mat-
ter of seconds. The temperature increments

CARNEGIE INSTITUTION

Fig. 20. Microphotograph of heater with glass
sample in place as indicated.

reported here were accomplished within
10-30 seconds. After a sample is melted
into the hole it is held in place by surface
tension during high temperature experi-
ments. Sample thickness near the edge is
about 500 Jim, and thicker in the center.
Although the emf from the thermocouple
yields an approximate temperature (uncer-
tainties are introduced by electrical con-
nections between dissimilar metals and al-
loys), accurate temperature calibration is
achieved by suspending a 0.04 mm Pt-
Pt9oRhio thermocouple into the melt. The
temperatures measured with this design are
accurate to within 4 °C anywhere within the
sample. The microheater was then placed
on the microscope stage of a custom de-
signed microscope port of a Dilor XY con-
focal micro-Raman system equipped with
an EG&G Model 1433-C cryogenic CCD
detector. Small chips (~ 1 mg) of the glass
were placed over the hole in the microheater
(Fig. 20) and melted. The samples were
excited with the 488-nm line of a Spectra
Physics model 2025 Ar+ ion laser, operat-

ing near 850 mW at the sample.

The compositions2 chosen for high-tem-
perature spectroscopy were on binary M2O-
Si02 joins, with M = Li, Na, andK. These
three different metal cations were chosen
so as to evaluate the influence of electronic
properties of the alkali metal on the tem-
perature-dependence of their structure.
Relationships between bulk melt polymer-
ization (NBO/Si) and the temperature de-
pendence of the structure were addressed
by changing the Si/O in each binary sys-
tem.

The widest range of metal oxide/silicon
compositions were studied in the system
Na20-Si02 (Fig. 21) as the problem of
glass devitrification appeared to be mini-
mal in this system. The high-frequency
region (800-1300 cnr*) of the NS7, NS5,
NS3, and NS2 glass spectra (marked 25 in
Fig. 21a, b, c, d) are similar to those re-
ported by Mysen et al. (1982). There is an
approximately 1 00 cm- 1 wide (full width at
half height) band centered near 1 100 cm-1.
A high-frequency shoulder near 1150
cm- * can be discerned in the spectra of NS7
and NS5 glasses. All spectra exhibit a dis-
tinct band near 950 cm-1, the intensity of
which visually increases systematically
with increasing Na/Si (increasing NBO/Si
of the melt). The glass spectra have been
deconvoluted previously (Mysen et al.,

2 These compositions are designated by the molar
ratio of M2O to Si02, so that NS7, NS5, NS3, and
NS2 are compositions Na207Si02, Na20»5Si02,
Na20*3Si02, and Na20«2Si02, and KS5, KS4,
LS3, and LS2 are K20»5Si02, K20»4Si02,
Li2O3Si02, Li20»2Si02, respectively.

GEOPHYSICAL LABORATORY

6^
CO

200

150

100 V

.'4

r* •: # - v.- .;%

fitist <

NS7

T|iq:1367°C
Tn :820°C

200

900 1000 1100 1200

Wavenumber, cm"

150

100

.*** V
WAS

NS5

liq

1255°C
945°C

^■N. ^ ^y v*v .*>.>.% vs «*. ■-< y ■.* t

•. •x::-;..::;;::v ..;**■ V-. .- *>. •

800 900 1000 1100 1200

Wavenumber, cm"1

1300

200 -

150 -

NS3

T„q: 805°C

Tg: 445°C

100

800 900 1000 1100 1200

Wavenumber, cm"

1300

200 -

150

100

VX ff£f*jJST

NS2

T,iq: 870'C

Tg: 490 °C

l\%\... ...... w,..%v1 1 55

900 1000 1100

Wavenumber, cm"

Fig. 21. In-situ, high-temperature Raman spectra of glasses, supercooled melts, and melts on the join
Na20-Si02 as a function of temperature (numbers on the right side of each spectrum represent
temperature in °C). The liquidus, Tug, and glass transition, To, temperatures for each composition are
indicated on the figures. The intensities are calculated relative to the greatest intensity within each
spectrum. A - composition NS7 (Na20»7Si02, bulk melt NBO/Si=0.28); B - NS5 (Na20»5Si02, bulk
melt NBO/Si=0.4); C - NS3 (Na203Si02, bulk melt NBO/Si=0.67); D - NS2 (Na202Si02, bulk melt
NBO/Si=l). In each panel, the spectra from successively higher temperature are offset by 10% for
clarity.

CARNEGIE INSTITUTION

100

'55

S 40

800

NS2

Na20-Si02 glass
25°C

[usa\ Jz

; >NS5 Nst.
* .?.,*«*{«'. *
* ,<NS7*

"ft L_

c
c

900

1000

1100

1200

1300

Wavenumber, cm"

100

A Na20-Si02 melts
*\| 1144-1165°C

• w

NS2 • ;

•

- #% . *

.•; :*

• .i

*. »*

• •

• if

■ •

• ■ •

. #•

• ••

# : * -•

9 \
• .i

r : V7

• . t

.-. •

•%

« .„-.

• >

* NS3 i .

• •:••

- • >S /:

*. **

; ; rHE&*$ •'

.; f :J*v^- •

t *

.. ' f JT • • .'

\ 'i

•fr' ?*\ NS7 /

vs.

800

900

1000

1100

1200

1300

Wavenumber, cm"

Fig. 22. Comparison of room temperature Raman spectra of glasses on the join Na20-Si02 (A) with
spectra of the same materials at temperatures between 1144°C and 1165°C (B) for the sample
compositions as shown in Fig. 21. The temperatures for all but the NS7 samples are above the liquidus
temperature. For the NS7 sample, the temperature of spectra acquisition is above the glass transition
temperature (see Fig. 21 for glass and liquidus temperatures of these samples). The intensities are
calculated relative to the greatest intensity within each spectrum.

1982) and the 950, 1100 and 1150 cm-1
bands have been assigned to Si-0 stretch
vibrations in specific structural units (see
Virgo et al, 1980; Furukawa et ai, 1981;
McMillan, 1984, for discussion of band
assignments). The sharp band near 950
cm-1 is assigned to Si-O" stretching in
structural units with 2 nonbridging oxygens
per silicon (Q2). The main band centered
near 1 100 cm-1 is due to Si-O" stretching in
units with NBO/Si = 1 (1100 cm1 band;

3 For convenience, in this paper, we will refer to
the 950, 1 100, and 1 150 cm"1 bands as the Q2, iP
and Q^ bands

Q3). The high frequency shoulder on the
1 100 cm- * band at approximately 1 1 50 cnr
1 results from the presence of fully poly-
merized units (Q4). Thus, all the glass spec-
tra from the Na20-Si02 system are consis-
tent with the existence of structural units
with their individual NBO/Si values = 2
(Q2 or Si032" units), 1 (Q3 or Si20s2"
units) and 0 (Q4 or Si02 units)3. There is no
evidence for structural units less polymer-
ized than NBO/Si = 2 in these spectra, in
accord with previous Raman and NMR
data (e.g., Virgo et al. 1980; My sen et al,
1982; Stebbins, 1987). Equation (1) can be

100 i_

:t»

0 #

NS5

NS5 & KS5
25°C

%•

'55

1100

1200

1300

Wavenumber, cm"

100

60 -

« # „ •

NS5 & KS5
1 370-1 380°C

• •

•

T^5:1255°C
Tf 5: 745°C

T^5: 940°C

TgS5: 540°C

•

"»•

•

» •

6 *

v #

'" •

, •

^ *

"•

" »

.3= *

r. •

o •

NS5 »w ;

»J

*4*.Vv .#

*_v

"e KS5 *

o tft&ffW'

i ■. •* i i

v&f20»K

800 900 1000 1100 1200

Wavenumber, cm'

1300

Fig. 23. Comparison of room temperature (25°C - A) and high temperature (1370-1380°C - B) spectra
of KS5 (K205SiC>2) and NS5 (Na20»5Si02) glasses (A) and melts (B). The bulk melt NBO/Si for both
samples is 0.4. The intensities are calculated relative to the greatest intensity within each spectrum.

used, therefore, to describe the anionic
equilibria for these compositions.

Increasing temperature has three effects
on the Raman spectra. (1) The band near
950 cm-1 (Q2 band) relative to the band
near 1100 cm-l increases in intensity as a
function of increasing temperature (Fig.
21 A through 2 ID). An increase of about
60% from the 25° C spectrum to the high-
est-temperature spectrum was determined
for all four compositions (Fig. 22). A simi-
lar increase in intensity with temperature
was also noted by Seifert tt al. ( 1 98 1 ) in the
high-temperature spectra of (Na2Si205)85-
(Na2(NaAl)205)i5 melt. (2) The intense
envelope centered near 1100 cm-1 broad-

ens with temperature (from slightly less
than 100 cm-1 in spectra of room tempera-
ture glass to more than 1 20- 1 30 cm- 1 in the
spectra of melts and supercooled liquids.
The -1150 cm-1 shoulder evident in the
NS5 and NS7 room temperature spectra is
not so clear. (3) The frequency of the -950
cm-1 band as well as the envelope centered
near 1100 cm-1, shifts to slightly lower
temperature.

The relationships between temperature
and electronic properties of the network-
modifying metal cation (K, Na, and Li) are
illustrated in Figs. 23 and 25. At room
temperature (Fig. 23A), the spectrum of
NS5 glass shows a more distinctive Q2

CARNEGIE INSTITUTION

1001-

NS3 & LS3
25°C

100

1 40

1300

/"A Nf

. * • 1 1

» : 1 •

NS3 & LS3
55-1165°C

T^S3: 805°C

LS3 /

NS3

g •

LS3

iiq:

»:T9

LS3

445 °C
1224°C
725°C

/ ns3 / ;\

QEteJL ■ i i ^^i

Wavenumber, cm

800 900 1000 1100

Wavenumber, cm

1200

-1

1300

100 1-

80 -

■*— •
CO

60 -

40 -

20;

800

rV. NS2 & LS2

9
•
r
•

:.: 25'C

• •

7
•
•

: »•
••

/ .'

••

••
•

,Als2,' :

••
••
••

% * 1

— •

•

ft <

•

\ j

•7

*NS2/

•
9

I t 1

I I l

100

I 40

900 1000 1100

Wavenumber, cm

1200
-1

1300

NS2 & LS2
1205-1 21 5°C

LS2," J

X / /

* 1/ /

? 'J

: Vns2

NS2
Iiq •

NS2.

g :

LS2.
Iiq ■

LS2

870°C
490*C
1033°C
600°C

_ ^*^^j

800 900 1000 1100

Wavenumber, cm

1200
-1

1300

Fig. 24. Comparison of Raman spectra of glasses and melts in the systems Li20-Si02 and
Na20 for trisilicate (bulk melt NBO/Si = 0.67) (A - room temperature and B - 1 1 55- 1 1 65°C)
and and disilicate (bulk melt NBO/Si = 1 .0) compositions (C - room temperature, D - 1 205-
1215°C). Symbols: NS3 - Na20«3Si02, LS3 - Li20»3Si02, NS2 - Na20«2Si02, LS2 -
Li20»2Si02. Relevant liquidus and glass transition temperatures are indicated on the
figures. The intensities are calculated relative to the greatest intensity within each spectrum.

band than that of KS5 (in both glasses, the
bulk melt NBO/Si = 0.4). When comparing
the spectra of LS3 and NS3 (NBO/Si =
0.667) and LS2 and NS2 (NBO/Si = 0.14)
(Figs. 24A and 24C), the intensity of the Q2
band is again more pronounced in the spec-

tra of glasses with the smallest metal cation
(Li). This increased intensity of the 950
cm-1 band would indicate enhanced abun-
dance of Si032_ (or Q2) structural units.
High -temperature spectra of KS5 and
KS4 melts reveal the same broadening of

GEOPHYSICAL LABORATORY

250 i-

200 -

150

I 100

KS5

T,iq:940°C
Tn : 536° C

900 1000 1100

Waven umber, cm

1200

-1

1300

250 -

200 -

KS4

Tliq: 772°C
: 423°C

& 150 P

— 100

800 900 1000 1100

Wavenumber, cm

1200

1300

Fig. 25. Raman spectra of KS5 (K20»5SiC>2) glasses and melts (A) and KS4 (K20«4SiC>2) glasses and
melts (B) as a function of temperature (°C) as indicated on the right side of each spectrum. The intensities
are calculated relative to the greatest intensity within each spectrum, and spectra at successively higher
temperatures are offset by 10 % for clarity.

the 1100 cm-1 envelope in the KS5 and
NS5 spectra (Fig. 25). The KS5 spectrum
does not show evidence for a 950 cm-1
band at any temperature studied, whereas
that of NS5 does with its intensity
increasing with temperature. In spectra of
the KS4 composition, the 950 cm-1 band is
present in the room temperature spectra as
a very weak band or shoulder in the glass
spectrum, and shows a distinctive intensity
increase with increasing temperature (Fig.
25B). Thus, it would appear that Si032_
structural units are no longer discernible
(within the sensitivity of the spectroscopic
technique) in glasses and melts on the K2O-
Si02 join for compositions with bulk melt
NBO/Si of 0.4. Increased temperature (at
least to 1 380°C) does not change this obser-
vation. In contrast, in the Na20-Si02 sys-
tem, all structural units appear to be present
at all temperatures with melts at least as

polymerized as NS7 (NBO/Si = 0.28).

The comparison of the glass and melt
spectra of the tri- and di-silicates of Na and
Li (Fig. 24) reveal (1) that the intensity of
the Q2 band in LS3 glass is greater by a
factor of about 3 compared with that of
NS3 (Fig. 24 A), whereas at high tempera-
ture in the molten range, the difference has
decreased to about a factor of 2 (Fig. 24B).
These relative intensity changes can also
been seen in the glass and melt spectra of
LS2 and NS2 composition (Fig. 24C,D).
(2) In the 1100 cm-1 envelope, the maxi-
mum is at lower frequency in the spectra of
the lithium samples than in the sodium
samples and there might be a slight in-
crease in frequency difference as the the
glasses are transformed to melts.

Previous deconvolutions of this high-
frequency envelope (e.g., My sen et al.,
1982; Mysen, 1990) demonstrated that the

CARNEGIE INSTITUTION

1150 cm-1 band (Si02 or Q4) on the high-
frequency limb of the 1100 cm-1 band
always increased when the intensity of the
Q2 band increased. These increases were
accompanied by a concomitant decrease in
the 1100cm-1 (Si2052" or Q3) band inten-
sity. Observations such as these, also con-
sistent with interpretation of 29Si NMR
spectra (e.g., Stebbins, 1987), lead to the
conclusion that in metal oxide silicate
glasses whose anionic equilibrium can be
described with equation ( 1 ), increased abun-
dance of Q2 (or Si032_) structural units is
always accompanied with an increase in
Q4 (Si02) and a decrease in Ql (Si2052)
units. The spectra of glass, supercooled
liquid, and liquid for Na20-Si02 and K2O-
Si02 show visual evidence for a shift of
equation (1) to the right with increasing
temperature. In the absence of detailed
statistical deconvolution, the evidence for
the glasses and melts in the Li20-Si02
glasses and melts is less obvious. It would
appear, from comparison of NS3 with LS3
and of NS2 with LS2 glass and melt spectra
that increasing temperature has less effect
on the spectra of Li-silicates than on those
of the Na-silicates. Visually, the intensity
near 950 cm-1 in the Li-silicate spectra
does not change significantly with tem-
perature, whereas those of NS3 and NS2
do. The intensity difference between the
two sets of spectra decreases, therefore, as
the glasses are heated and eventually melted.
In summary, high-quality Raman
spectra of silicate melts can be recorded in-
situ at magmatic temperatures and above
with sample acquisition times on the order
of one minute or less. From such spectra, it
has been found that in binary alkali metal-

silica systems in the bulk melt polymeriza-
tion range between 0.28 and 1.0, glasses,
supercooled melts, and melts in the tem-
perature range 25-1475°C generally con-
sist of coexisting Si032" (g2), Si2052"
(Q3), and Si02 (Q4) structural units. In
potassium-bearing systems, the upper NBO/
Si limit for Si032~ units probably is at
NBO/Si between 0.4 and 0.5. No tempera-
ture-dependence of this limit was observed.
No additional units were identified within
this temperature range. For compositions
with the Z/r2 of the alkali metal ranging
from 2.8 to 0.6 (Li, Na, and K), equation (1)
shifts to the right with increasing tempera-
ture. The spectra probably indicate a com-
positional dependence of the free energy
for reaction (1). Qualitatively, the effect of
temperature on equilibrium (1) decreases
as the Zjr2 of the metal cation increases
(Li>Na>K) , but quantitative evaluations of
its values have not been been carried out.

References

Brandriss, M. E., and J. F. Stebbins, Effects of
temperature on the structures of silicate liq-
uids: 29Si NMR results, Geochim. Cosmochim.
Acta, 52, 2659-2669, 1988.

Cooney, T. F., and S. K.Sharma, High temperature
Raman spectral study of Ge02 andRb4SigOi8
crystals, glasses and melts, EOS, 71, 1672,
1990.

Furukawa, T., K. E. Fox, and W. B. White, Raman
spectroscopic investigation of the structure of
silicate glasses. HI. Raman intensities and
structural units in sodium silicate glasses, J.
Chem.Phys., 153, 3226-3237,1981.

Matson, D. W., S. K. Sharma,and J. A. Philpotts,
The structure of high-silica alkali-silicate
glasses — A Raman spectroscopic investiga-
tion, J. Non-Cryst. Solids, 58, 323-352, 1983.

McMillan, P., A Raman spectroscopic study of
glasses in the system CaO-MgO-Si02, Amer.
Mineral., 69, 645-659, 1984.

GEOPHYSICAL LABORATORY

Mysen, B. O., The role of aluminum in depoly-
merized, peralkaline aluminosilicate melts in
the systems Li20 - AI2O3 - Si02, Na20-
Al203-Si02 and K20-Al203-Si02, Amer.
Mineral., 75, 120-134, 1990.

Mysen, B. O., D. Virgo, and F. A. Seifert, The
structure of silicate melts: Implications for
chemical and physical properties of natural
magma, Rev. Geophys., 20, 353-383, 1882.

Seifert, F. A., B. O. Mysen, and D. Virgo, Struc-
tural similarity between glasses and melts rel-
evant to penological processes, Geochim.
Cosmochim. Acta, 45, 1879-1884, 1981.

Stebbins, J. F., Effects of temperature and compo-
sition on silicate glass structure and dynamics:
Si-29 NMR results, /. Non-Cryst. Solids, 106,
359-369, 1988.

Stebbins, J. F., Identification of multiple structural
species in silicate glasses by 29Si NMR, Na-
ture, 330, 465-467, 1987.

Virgo, D., B. O. Mysen,and I. Kushiro, Anionic
constitution of silicate melts quenched at 1 atm
from Raman spectroscopy: Implications for
the structure of igneous melts, Science, 208,
1371-1373, 1980,

Peralkalinity and H2O Solubility
Mechanisms in Silicate melts

Bjorn Mysen

Relationships between properties of
magmatic liquids and their water content
have been examined since the pioneering
work of Bowen (1928). Despite an exten-
sive literature on the subject, the detailed
nature of the interaction between H2O and
the silicate melt structure remains unclear,
and characterization of the relations be-
tween the structure of hydrous silicate melts
and their physical and chemical properties
remains a topic of intense interest.

The principal anionic equilibrium that
describes the melt structure in magmatic
liquids under anhydrous conditions is (Virgo
etal, 1980; Mysen et al., 1982; Matsonef
al, 1983; Stebbins, 1987)

T2O52- <^> TO32- + IO2.

(1)

Dissolved H2O can interact not only with
TO2. but with all the structural units in the
melts. A study has been conducted, there-
fore, under conditions of constant degree
of bulk melt polymerization (NBO/T of
anhydrous melts is 0.5) with Al/(A1+Si)
and water content as the compositional
variables. The NBO/T value corresponds to
magma compositions intermediate between
tholeiite and andesite (Mysen, 1988). The
range in Al/(A1+Si) (0-0.3) covers that found
in most natural magmatic liquids.

Starting materials were mixtures of
Na2C03+Al203 + Si02 on thetetrasilicate
(Na2Si409)-tetra-aluminate
[Na2(NaAl)409] join (Fig. 26). Exchange

Fig. 26. Composition of anhydrous starting mate-
rials superimposed on simplified liquidus phase
relations in the system Na20-Al203-Si02 (from
Osborn and Muan, 1960).

CARNEGIE INSTITUTION

Anhydrous

850 1075 1300

Wavenumber, cm"1

- 100r

7.5wt%H20

830 1065 1300

Wavenumber, cm-1

Fig. 27. Curve-fitted Raman spectra of composi-
tions indicated. The V950 and vnoo bands are
shaded for clarity.

of Al3+ for Si4+ does not affect the bulk
melt polymerization (NBOIT = 0.5) under
anhydrous conditions because both Al3+
and Si4+ occupy tetrahedral coordination
(Mysen, 1990).

The glass starting materials were from
the same batch of glasses used by Mysen
(1990). About 20 mg of finely crushed
glass together with distilled, deionized H2O
was placed in sealed Pt containers for high-
pressure synthesis. All water contents (<7.5
wt %) were less than that needed to saturate
the melts atthe 12kbarand 1400°Cusedfor
sample preparation. These samples were
subjected to 12 kbar at 1400°C in the solid-
media, high-pressure apparatus (Boyd and
England, 1960) for 90 min and tempera-
ture-quenched at a quenching rate near
100°C/s between the experimental tem-
perature and ~500°C. The quenching rates
were similar for all materials studied. The
pressure uncertainty (as calibrated against
the quartz <=> coesite and albite <=> jadeite +
quartz transitions) is near ±1 kbar. The
effect of pressure on the Pt-Pt9oRhio ther-
mocouples is about ±10°C (Mao etal., Year
Book 70, p. 281-287). With no pressure
correction on the emf output from this
thermocouple, the temperature is consid-
ered accurate to ±10°C.

Structural information was obtained
from analysis of Raman spectra of the
quenched melts. The spectra were recorded
with an automated single-channel Raman
spectrometer system with the frequency-
doubled 532-nm line of an NdYAG laser
operating at 1 W for sample excitation.

The abundance of the structural units
was determined from the Raman spectra,
as described by Mysen {Annual Report

GEOPHYSICAL LABORATORY

1988-1989, p. 47-54). From deconvoluted,
high-frequency Raman spectral envelopes
such as illustrated with Figure 27, the rela-
tive intensities of the V950 [(Si, Al)-0" stretch
band from TO32- structural units] and vi 100
bands (Si, Al)-0 stretch band from T2O52-
structural units] were used for this purpose.
The distribution of Al among those units as
well as the fully polymerized structural
unit (7T)2.; V1150 and V1200 bands) was
evaluated from frequency shifts of the indi-
vidual bands as a function of H2O content
and bulk melt Al/(A1+Si).

There are systematic shifts in frequency
of important Raman bands with increasing
water content and with increasing bulk Al/
(Al+Si) (Fig. 28). The general descent of
the vi 150 and V1200 frequencies with in-
creasing H2O, also noted for hydrous
NaA102-Si02 melts (Mysen and Virgo,
1986), indicates that Al/(A1+Si) of the
fully polymerized structural units (TO2.) in
the melts is positively correlated with H2O
concentration. It is notable, however, that
even in anhydrous samples, the V1150 and
vi 200 frequencies at given Al/( Al+Si) are
lower than those observed in the spectra of
fully polymerized (NBO/T = 0) NaA102-
Si02 quenched melts with the same bulk
Al/(A1+Si) (Fig. 28). If the frequency of
these bands were used as a quantitative
measure of Al/( Al+Si) in the 7X)2. units,
[Al/Al+Si)]7T)2. in the anhydrous sodium
aluminosilicate melts would be 0.32, 0.34,
and 0.37, for NS4-A10 [bulk melt Al/
(Al+Si) = 0.1], NS4-A20 [bulk melt Al/
(Al+Si) = 0.2], and NS4-NA30 [bulk melt
Al/( Al+Si) = 0.30], respectively. The er-
rors in these numbers, however, are quite
large due to the shallow slope of the fre-

1300

§ 1200

900

Anhydrous

_» 1 i-

0.0 0.1 0.2 0.3 0.4 0.5 0.6
AI/(AI+Si)

1300

> - »

o 1200CT--

tf^- - -O

§ 1100,
ex

1000

3 wt% H2O

900

0.1 0.2 .0.3 0.4
AI/(AI+Si)

1300

1000

900

5wt%H20

0.0 0.1 0.2 0.3 0.4
AI/(AI+Si)

Fig. 28. Frequencies of selected (indicated on
Figure) Raman bands as a function of bulk melt
Al/(A1+Si) and H20.The bands shown are as-
signed to (Si,Al)-0° stretch vibrations in fully
polymerized structural units. The dashed line is
from the system NaA102-Si02-H20 (data from
Mysen and Virgo, 1986).

CARNEGIE INSTITUTION

1.0
0.8

NS4;AI/(AI+Si)=0; B r NS4-A10; Al/(AI+Si)=0.

6 0 2

wt% H20

Fig. 29. Abundance of structural units (as shown on figure) as a function of H2O content
and Al/(A1+Si). Open symbols are results from samples with D2O.

quency vs. Al/(A1+Si) curves. The results
are, nevertheless, consistent with a distinct
preference of Al3+ for the most polymer-
ized among the coexisting structural units
in the anhydrous melts, a conclusion con-
sistent with other Raman and NMR data
(e.g., Mysen etal, 1981; Engelhardt etal.,
1985; Kirkpatrick et al., 1986; Oestrike et
al., 1987; Domine and Piriou, 1986;
Merzbacher et al, 1990; Mysen, 1990).
Thus, the Al/(A1+Si) of the TO2 structural
units would exceed that of the bulk melt.
The frequency reduction of the V1150
and vi 200 bands with increasing H2O con-
tent leads to the suggestion that the [Al/
Al+Si)]7U2 in hydrous melts is enhanced

further. The effect of water on these Raman
frequencies is less pronounced, however,
in the peralkaline aluminosilicate melts
studied here than in hydrous NaA102-Si02
melts (Mysen and Virgo, 1986). If the fre-
quency trajectories in spectra from anhy-
drous Si02-NaA102 quenched melt were
employed to calibrate the Al/(A1+Si) of the
fully polymerized units in hydrous melts,
the minimum frequency (near 1080 cm-1)
corresponds to Al/(A1+Si) > 0.4 for for
7T)2. units hydrous aluminosilicate melts.
The influence of water on the abun-
dance of structural units in the melts (Fig.
29) can be inferred from the intensity varia-
tions of the bands. At low water contents,

GEOPHYSICAL LABORATORY

the AT2O5 in hydrous NS4 increases (the
melts become depolymerized) with water
content. This hydroxy lation mechanism
may be expressed with the equation

2Si02 + H20 » SiO(OH)2,

(2)

which when considered with equation (1)
probably results in an increased abundance
of S12O52- accompanied by a similar de-
crease in Si032_. If this were the only
solubility mechanism, the rate of depoly-
merization would be 0.0133 per mol %
dissolved H2O.

With high H2O contents there is evi-
dence, however, that these melts undergo a
slight polymerization (Fig. 29 A). Hydroxy -
lation of network-modifying Na+ can be
described by the system of equations

2Na+ + H20 + Si032-

<=> 2Na(OH) + Si02,

2Na+ + H20 + Si2052"

«> 2Na(OH) + 2Si02,

and

2Na+ + H20 + 2Si032"

<=> 2Na(OH) + Si2052-,

(3)

(4)

(5)

in which the network-modifying cation
(Na+) reacts with water to form Na(OH).
Thus, even in this compositionally
simple Na20-Si02-H20 system, the spec-
tral data are consistent with water solubil-
ity mechanisms that include three different
types of OH-bonding [H-OH (molecular
H2O), Na-OH (to form NaOH complexes),
and Si-OH.

The abundance trends of structural units
in the aluminous samples have four fea-
tures in common. (1) Whether anhydrous
or water-bearing, the abundance of fully
polymerized structural units is several tens
of percent higher than in the absence of Al
(Figs. 29B-D). (2) The abundance of TO2
passes through a maximum with water con-
tents between 1.5 wt %, and 3 wt %. (3)
Both depolymerized ^Os2" (NBO/T= 1)
andTC>32 (NBO/T=2) units generally coex-
ist with fully polymerized TO2 (NBO/T=0).
(4) There is a minimum in abundance of
depolymerized units (T2O5 and TO3) at
water contents corresponding to the maxi-
mum TO2 concentration.

The abundance patterns of the struc-
tural units in Al-bearing samples indicate
that at low H2O concentration (<1.5 wt %
H2O) solution of water results in polymer-
ization of the melts. The solubility mecha-
nism consistent with polymerization is in-
teraction between Na+ and H2O along the
lines of equation (3).

The Al-bearing melts become depoly-
merized with H2O contents > 1 .5-3 wt % as
the abundance of TO3 and T2O5 units is
positively correlated with H2O concentra-
tion, whereas that of 702. is negatively
correlated (Fig. 29B-D). Depolymeriza-
tion via interaction of H2O with expulsion
of tetrahedrally coordinated Al3+ (because
of reduction in Na+ charge -balance, or
Al(OH)3 formation, or both) is a principal
structural mechanism describing this be-
havior.

Although Al3+ does not reside exclu-
sively in fully polymerized anionic units
[the frequencies of the V95oand vi 100 bands
are weakly dependent on Al/(A1+Si), in

CARNEGIE INSTITUTION

particular in the low water concentration
ranges], available information (e.g,
Engelhardt et al., 1985; Kirkpatrick et al.,
1986; Oestrike et al, 1987; Mysen, 1990)
indicates a strong preference of aluminum
for the most polymerized of the coexisting
units at least in anhydrous melt systems.
We will discuss, therefore, the solubility
mechanism on the basis of all Al in such
fully polymerized units and refer to them as
NaA102. The principles derived from that
discussion will not, however, be affected
by some Al3+ in more depolymerized struc-
tural units.

The depolymerization reactions of wa-
ter-bearing alkali aluminosilicate melts rest
on the premise that if Al3+ in 4-fold coordi-
nation in anhydrous melts forms Al(OH)3
complexes after hydrolysis (with Al no
longer in tetrahedral coordination), a frac-
tion of the charge -balancing cation (Na+ in
the present system) equal to the proportion
of Al3+ in these complexes becomes net-
work-modifying. This mechanism could
be described with the following set of
equations:

2NaA102 + 2Si02 + 3H20

<^> 2A1(0H)3 + 2Na+ + Si2052", (6)

2NaA102 + Si02 + 3H20

^> 2A1(0H)3 + 2Na+ + Si032", (7)

2NaA102 + Si2052" + 3H20

<=> 2A1(0H)3 + 2Na+ + 2Si032". (8)

An alternative or additional mechanism
would operate if the charge -balancing cat-
ion (Na+) interacts with dissolved H2O to
form OH complexes. Then, an equivalent

fraction of Al3+ originally in tetrahedral
coordination will become a network-modi-
fying cation.

2NaA102 + 3Si02 + H20

<=> 2Na(OH) + 2A13+ + Si032', (9)

2NaA102 + 6Si02 + H2O

<=> 2Na(OH) + 2A13+ + 3Si2052-, (10)

2NaA102 + 3Si2052" + H20

<=> 2Na(OH) + 2A13+ + 6Si032". (11)

Manning et al. (1980) and Pichavant
(1987) inferred from liquidus phase equi-
libria in hydrous quartz-feldspar systems
that this mechanism would explain their
observations. Kohnera/. (1989) interpreted
their NMR spectra of hydrous NaAlSi30s
glass as consistent with NaOH complexing.
In the case of Na-aluminosilicate melts, the
latter mechanism (NaOH complexing) more
efficiently depolymerizes the melt, be-
cause for each charge-balancing Na+ in
anhydrous melts that forms the NaOH com-
plex in the hydrous environment, three
nonbridging oxygens can be stabilized with
Al3+. For each tetrahedrally coordinated
Al3+ in anhydrous melts, transformed to
Al(OH)3 in hydrous melts only one
nonbridging oxygen can be stabilized with
the Na+ released in this process.

These expressions are consistent with
the observations in Figs. 29B-D, where the
abundance of both T2O52- and TO32
increasesand that of TO2 decreases as the
water concentration increases above 1.5
wt%. Equation (5) (depolymerization) to-
gether with equation (3) (polymerization)
may be an adequate approximation to the

GEOPHYSICAL LABORATORY

principal solution mechanism of H2O in
highly aluminous, melts such as NS4-A30
[Al/(A1+Si)=0.3].

Liquidus phase relations in the system
Na20-Al203-Si02 in the compositional
region near the Na2Si409-Na2(NaAl)409
join can be employed to illuminate the
relations between H2O activity and liquidus
phase relations. Due to the lack of experi-
mental data on necessary high-pressure
liquidus phase relations, the 1 -bar informa-
tion (Osborn and Muan, 1960) will be used
for the purpose. It is recognized that by
using the 1-bar data, significant errors are
introduced, as it is well established that
pressure by itself affects both liquidus tem-
peratures and liquidus volumes in this (and
other) systems (e.g., Boettcherera/., 1984).
Nevertheless, it is informative to discuss
the consequences of dissolved water only.
Isopleths of 2.0 and 7.5 wt % were recast
in terms of Na20, AI2O3, and Si02. The

liquidus temperatures were corrected for
water content by assuming ideal mixing of
H2O in melts on the basis of the 8 -oxygen
model of Burnham (1975) and by using the
heat of fusion data for liquidus phases
summarized by Richet and Bottinga ( 1 986)
with the expression

7(K)

-1— fail

v melt\

AHoo/

(12)

a o fusion

where it was assumed that the heat of
fusion of the relevant phases was tempera-
ture independent.

It is evident that with small amounts of
water in melt solution (2 wt % H2O), the
albite liquidus volume is greatly expanded
at the expense of both quartz on its low-Al
side and nepheline on its high- Al side (Fig.
30). The small liquidus surface of crystal-
line sodium disilicate (NS2; Na2Si20s) in

1200

1000

£ 800

■*— »
co

g. 600

E
a>

•" 400

200

■ $\

— 1

^ Ab

1 T ■

""'Ab y
NS2

— , 1 1 1

1 _ ^ 1

^Ab
1

NS2

•

0.0

0.1

0.2
AI/(AI+Si)

0.3

0.4

Fig. 30. Calculated liquidus surfaces along the anhydrous (NBO/T = 0.5), 2 and
7.5 wt % H20 isopleths, as a function of Al/(A1+Si).

CARNEGIE INSTITUTION

the anhydrous system has disappeared. Ex-
pansion of the albite liquidus volume at low
PH20 (-1.25 kbar) along the join
NaAlSi308-Na2Si205 has also been docu-
mented by experimental phase equilibrium
results (Mustart, 1972). The liquidus sur-
faces of fully polymerized albite and neph-
eline and complete lack of depolymerized
crystalline phases reflect, therefore, the
tendency toward polymerization of the
peralkaline aluminosilicate melts with small
amounts of water in solution.

With 7.5 wt % H2O in solution the
liquidus temperatures are greatly depressed
(Fig. 30). There is no quartz liquidus vol-
ume, and crystalline NS2 (sodium disilicate)
is an important phase from Al-free compo-
sitions to Al/(A1+Si) near 0.15, where
albite appears. The albite liquidus volume
ranges from Al/(A1+Si) = 0.15 to about
0.22, whereas with 2 wt % H2O the volume
ranges from near 0.05 to about 0.26. This
very significant expansion of the NS2
liquidus volume at the expense of albite is
a consequence of the depolymerization of
the melt caused by the dissolved water. The
activity of the nepheline component is in-
creased relative to that of albite as reflected
in the encroachment of the nepheline
liquidus volume on that of albite. Even
though the abundance of TO2 structural
units has been lowered, most likely this
expansion of nepheline relative to albite
results from an enhancement of Al/(A1+Si)
in the remaining fully polymerized struc-
tural units.

References

Boettcher, A.L., Q, Guo, S. Bohlen, andB. Hanson,
Melting of feldspar-bearing systems to high
pressures and the structures of aluminosilicate
liquids, Geolog, 12, 202-204, 1984.
Bowen, N. L. The Evolution of the Igneous Rocks,
332 pp., Princeton Univ. Press, Princeton, 1 928
Boyd, F. R., and England J. L., Apparatus for
phase equilibrium measurements at pressures
up to 50 kilobars and temperatures up to
1750°C,/. Geophys.Res., 65, 741-748, 1960.
Burnham, C. W\, Thermodynamics of melting in
experimental silicate-volatile systems,
Geochim. Cosmochim. Acta, 39, 1077-1084,
1975.
Domine, R, and B. Piriou, Raman spectroscopic
study of the Si02-Al203-K20 vitreous sys-
tem: Distribution of silicon and second neigh-
bors, Amer. Mineral, 71, 38-50, 1986.
Engelhardt, G., M. Nofz, K. Forkel, F. G.
Wishmann, M. Magi, A. Samoson, and E.
Lippmaa, Structural studies of calcium alumi-
nosilicate glasses by high resolution solid state
29Si and ^Al magic angle spinning nuclear
magnetic resonance, Phys. Chem. Glasses, 26,
157-165, 1985.
Kirkpatrick, R. J., R. Oestrike, C. A. Weiss, K. A.
Smith, and E. Oldfield High-resolution 27A1
and 29Si NMR spectroscopy of glasses and
crystals along the join CaMgSi206-CaAl2Si06,
Amer. Mineral, 77,705-711,1986.
Kohn, S. C., R. Dupree, and M. E. Smith, A
Nuclear magnetic resonance study of the struc-
ture of hydrous albite glasses, Geochim.
Cosmochim. Acta, 53, 2925-2935, 1989.
Manning, D. A. C, C. M. B. Hamilton C. M. B.
Henderson, and M. J. Dempsey, The probable
occurrence of interstitial Al in hydrous F-
bearing and F-free aluminosilicate melts,
Contr. Mineral. Petrol, 75, 257-262, 1980.
Matson, D. W., S. K. Sharma, and J. A. Philpotts,
The structure of high-silica alkali-silicate
glasses — A Raman spectroscopic investiga-
tion, /. Non-Cryst. Solids, 58, 323-352, 1983.
Merzbacher, C, B. L.Sherriff, S. J. Hartman, and
W. B. White, A high-resolution 29Si and 27A1
NMR study of alkaline earth aluminosilicate
glasses, /. Non-Cryst. Solids, 124, 194-206,
1990.
Mustart, D. A. Phase relations in the peralkaline
portion of the systemNa20-Al203-Si02-H20 .
Ph. D. thesis, Stanford University , 1972.
Mysen, B. O., Effect of pressure, temperature, and
bulk composition on the structure and species
distribution in depolymerized alkali alumino-
silicate melts and quenched melts, J. Geophys.
Res., 95, 15733-15744, 1990.

GEOPHYSICAL LABORATORY

My sen, B. O., Structure and Properties of Silicate
Melts, 354 pp., Elsevier, Amsterdam, 1988.
Mysen, B. O., D. Virgo, and F. A. Seifert, The
structure of silicate melts: Implications for
chemical and physical properties of natural
magma, Rev. Geophys., 20, 353-383, 1982.
Mysen, B. O., and D. Virgo, The structure of melts
in the system Na20-CaO-Al2C>3-Si02-H20
quenched from high temperature at high pres-
sure. 2. Water in melts along the join NaAlC>2-
Si02 and a comparison of solubility mecha-
nisms of water and fluorine, Chem. GeoL, 57,
333-358, 1986.
Mysen, B. O., D. Virgo, and I. Kushiro, The
structural role of aluminum in silicate melts —
A Raman spectroscopic study at 1 atmosphere,
Amer. Mineral, 66,678-701,1981.
Oestrike, R., W.-H. Yang, R. J. Kirkpatrick, R.
Hervig, A. Navrotsky, and B. Montez, High-
resolution 23Na, 27A1 and 29Si NMR spectros-
copy of framework-aluminosilicate glasses,
Geochim. Cosmochim. Acta, 51, 2199-2210,
1987.
Osborn, E. F., and A. Muan, Phase equilibrium
diagrams for ceramists. Plate 4. The system
Na20-Al203-Si02, Am. Ceram. Soc., Colum-
bus Ohio, 1960.
Pichavant, M., Effects of B and H2O on liquidus
phase relations in the haplogranite system,
Amer. Mineral, 72,1056-170,1987.
Richet, P. , and Y. Bottinga, Thermochemical prop-
erties of silicate glasses and liquids: A review,
Rev. Geophys., 24, 1-26, 1986.
Stebbins, J. F.,. Identification of multiple struc-
tural species in silicate glasses by 29Si NMR,
Nature, 330, 465-467, 1987.
Virgo, D., B. O. Mysen, and I. Kushiro, Anionic
constitution of silicate melts quenched at 1 atm
from Raman spectroscopy: Implications for
the structure of igneous melts, Science, 208,
1371-1373.,1980.

Partitioning of Fluorine and Chlorine

between Apatite and Non-Silicate Fluids

at High Pressure and Temperature

James Brenan

mineral as a useful indicator of halogen
activities in the geologic environment. Prior
experimental work has focused on calibrat-
ing apatite as a monitor of halogen activi-
ties in the low pressure(f>)-temperature(7T)
hydrothermal environment (Korzhinskiy,
1981; Latil and Maury, 1977). Recently
developed solution models for apatite, flu-
ids and melts now provide a basis for theo-
retical prediction of the halogen chemistry
of apatite coexisting with these phases over
a somewhat broader range of P and T
(Candela, 1986;PiccoliandCandela, 1991;
Tacker and Stormer, 1989; Zhu and
Sverjensky, 1991). Although such work
represents an important contribution to the
interpretation of the halogen chemistry of
natural apatite, its usefulness may be lim-
ited to the overall low P-T range in which
experiments were performed for both cali-
bration and development of solution-model
databases. In order to exploit apatite as a
monitor of the halogen chemistry of fluids
or melts in the high P-T environment, ex-
perimental determination of apatite chem-
istry at these conditions is a requisite. This
report describes the results of experiments
aimed at this goal with specific emphasis
on measurements of the distribution of
fluorine and chlorine between apatite and
non-silicate fluids (H2O + dissolved salts ±
CO2, carbonate melt) at P-T conditions
appropriate to the lower crust and upper
mantle.

The presence of fluorine and chlorine as
essential structural constituents of apatite,
combined with the widespread occurrence
of apatite amongst diverse parageneses,
underscores the potential utility of this

Experimental Technique

Owing to the difficulties associated with
analyzing fluids and low-viscosity melts

CARNEGIE INSTITUTION

quenched from high P-T experiments, par-
tition coefficients (D-values; wt % concen-
tration in apatite/wt % concentration in
fluid or melt) were determined by mass
balance. The overall strategy was therefore
to perform experiments by (1) encapsulat-
ing known quantities of finely-powdered,
natural apatite (previously well-character-
ized in terms of fluorine and chlorine abun-
dances) and fluid, or apatite and melt, in
welded Pt containers, (2) equilibrating this
mixture for 2-4 days at high pressure and
temperature (950-1050°C, 1.0-2.0 GPa)
with a solid-media, high-pressure appara-
tus, then, (3) analyzing the run-product
apatite for F and CI using an electron mi-
croprobe. Fluids were added to these ex-
periments as distilled H2O or as aqueous
solutions of HC1 (1.8 and 5.3 wt %), NaCl
(4, 15, and 25 wt %), Na2CC>3 (4 and 15 wt
%) and NaOH (4 wt %). Water-carbon
dioxide mixtures were produced by weigh-
ing in distilled H2O with silver oxalate.
The carbonate melt used in these experi-
ments was a mixture of high-purity carbon-
ates with the stoichiometry 80 wt % dolo-
mite: 20 wt % Na2CCh. Water was added to
melt-bearing experiments to yield 5-15 wt
% in the melt. The apatite starting material
consisted of hand-picked fragments (free
of inclusions and visible alteration) of gem-
quality crystals that originate at Durango,
Mexico. Microprobe analysis of points on
25 different apatite fragments yielded aver-
age concentrations for F and CI of 3.57 (±
0.13, lo) and 0.38 (± 0.06, la) wt %,
respectively. A rastered beam was used to
minimize mineral degradation during mi-
croprobe analysis, and no loss of F or CI x-
ray intensity was detected even for data

acquisition times exceeding 4 minutes; stan-
dard analytical conditions were a 45 nA
sample current at 15 kV accelerating volt-
age.

Due to the low solubility of apatite in
C02-H20-NaCl and dilute HC1 solutions
(i.e.,<l wt%;Ayers and Watson, 1991), no
correction for apatite dissolution was ap-
plied to calculated partition coefficients
involving these compositions. Based on
the measurements of Ayers and Watson
(1991), a solubility of 3 wt % apatite was
used in partition coefficient calculations
for 5.3 wt % HCl-bearing experiments and,
combining the data of Baker and Wyllie
(1990) with that of Watson (1980), a solu-
bility of 12.5 wt % apatite was used to
calculate D-values for experiments involv-
ing carbonate melt.

100 1 — ■ — 1 — ■ — 1 — ■ — 1 — ■ — ' — ■ — > — ■ — ' — ■ — ' —

^ if-A - 2.2-2.6 Fluorme

'=3

.01

.001

*«* rwv

ft TA A AA

Chlorine

m D

03 n

2 4 6 8 10 12 14
Wt % CI

Fig. 31. Apatite/fluid partition coefficients for
fluorine (closed symbols) and chlorine (open sym-
bols) as a function of total chlorine abundance.
Data are from experiments at 2.0 GPa and 1050°C
and pertain to H20-bearing fluids with and with-
out added HC1 or NaCl (Circles, triangles and
squares refer to experiments with H2O, H2O-HCI
and H20-NaCl, respectively).) The numbered
contours for the fluorine partitioning data refer to
the total fluorine contents of these experiments
(rev = reversal).

GEOPHYSICAL LABORATORY

Results and Discussion

Inspection of apatite run products re-
vealed substantial grain-growth during the
course of an experiment: samples starting
out as <10 Jim, angular grains underwent
coarsening to produce subhedral to euhedral
material of from 15 jum to mpre than 100
Jim grain size. The amount of grain growth
was dependent on fluid composition and
experiments containing HC1 or Na-carbon-
ate solutions or carbonate melt produced
the most abundant large grains. Multiple
microprobe analyses across individual apa-
tite grains from run-products showed no
evidence of compositional zoning.

Figures 31 and 32 portray D-values for
fluorine and chlorine as functions of total
chlorine and fluorine concentration mea-
sured in experiments with H20-NaCl-HCl
fluids at 1050 °C and 2.0 GPa. (Note that F
concentrations were varied by adjusting
the ratio of apatite to fluid; CI contents
were varied by changing the CI concentra-
tion in added solutions.) Average/) -values
for CI from experiments containing H2O or

100

a. 1

.01

.001

Fluorine

rev

*•

Chlorine

*» A ^- reV

1
4&

*HCI

D O

2 NaCI

. 1

Oh2o

1 2

wt%F

Fig. 32. Apatite/fluid partition coefficients for
fluorine and chlorine as a function of total fluorine
abundance (symbols as in Fig.31). Data are from
the same experiments described in Fig. 31.

HCl-solutions are -0.1 and average values
obtained from aqueous NaCl-bearing ex-
periments are -0.015; no dependence on
either absolute F or CI concentration was
seen for either of these values. Consistent
with previous work (Korzhinskiy, 1981;
Latil and Maury, 1977), D-values for F are
well in excess of those measured for CI, and
are generally similar at a given total CI
concentration (for similar total F contents),
regardless of the mode of CI addition. Par-
tition coefficients for F exhibit no system-
atic dependence on total CI concentration
(Fig. 31), but values do systematically de-
crease as a function of increasing F concen-
tration (Fig. 32). Results of a reversal
experiment in which apatite was first equili-
brated in H2O, then in 1.8% HC1, gave
good agreement with D-values measured
in forward experiments. Average partition
coefficients for F and CI in basic solutions
(aqueous NaOH or Na2C03) are ~5 and
<0.02, respectively.

Figure 33 portrays fluorine and chlorine
D-values as a function of pressure for ex-
periments with aqueous HC1 or NaCI at
1050°C. D- values for CI are invariant with
pressure, regardless of fluid composition,
as is the D for F in the HCl-bearing fluid.
The F partition coefficient with aqueous
NaCI at 1.0 GPa is, however, -10 times
higher than values determined at 2.0 GPa
for similar total F abundances. D-values
for CI at 950°C and 2.0 GPa (Fig. 33) are
identical to values measured at 1050°C
whereas F partition coefficients are 4-10
times higher.

Fluorine and chlorine partition coeffi-
cients generally decrease and increase, re-
spectively, as a function of the mole frac-

CARNEGIE INSTITUTION

100

5= 1
0)

■«— »
TO

Fluorine

.01

.001

950°C

HCI

Chlorine
A —

950°C

1.0 1.5 2.0

Pressure (GPa)

2.5

Fig. 33. Apatite/fluid partition coefficients for
fluorine (closed symbols) and chlorine (open sym-
bols) as a function of pressure for experiments at
950 and 1050°C (1050°C data unlabelled). Data
are from experiments with 2.2-2.5 wt % total
fluorine (variable chlorine contents) with added
HCI (circles) or NaCl (triangles).

Hon of CO2 [X(C02>] in H2O-CO2 mix-
tures (Fig. 34). In terms of D-values for CI,
this relation is greatly altered for experi-
ments in which CI was added as solutions
with >15 wt % NaCl (Fig. 35; F partition
coefficients are relatively unaffected by

100

10 ■

■g

"=3

0) .
♦J

03
Q.

.01

1 ■

>
>

•

■

•

fluorine

chlorine

<D ;

)

■

— 1__

■

0.0

0.2 0.4 0.6

Mol Frac C02

0.8

Fig. 34. Apatite/fluid partition coefficients for
fluorine (closed symbols) and chlorine (open sym-
bols) as a function of the mole fraction of CO2 for
experiments involving CO2-H2O mixtures. Data
are from experiments run at 1050°C and 2.0 GPa.

i -1

O
Q

.01

H20

XC02 = 0.3-0.5
5% HCI

4% NaCl

15%NaCI
25% NaCl

1 2 3

Total CI (wt %)

Fig. 35. Apatite/fluid partition coefficients for
chlorine as a function of total chlorine concentra-
tion. Data are from experiments involving CO2-
H2O fluids with mole fractions of CO2 = 0.3-0.5.
Data labels refer to experiments with no added
chlorine (H2O; CI is from the apatite starting
material) or in which CI was added as 5 wt % HCI
or 4-25 wt % NaCl.

these fluid composition variations). As
illustrated in Fig. 35, similar D-values for
CI (i.e., -0.25) were obtained from experi-
ments in which CI was added as either the
apatite starting material, 4 wt % NaCl, or
5.3 wt % HCI. Experiments in which CI
was added as 1 5 or 25 wt % NaCl, however,
exhibit marked drops in the chlorine D-
value (i.e., to -0.065 and -0.03, respec-
tively). Inasmuch as CI partition coeffi-
cients were found to be independent of total
CI content for C02-free fluids, these results
may be somewhat surprising. The overall
reduction in the CI partition coefficient
could, however, be accounted for if an
H20-rich fluid evolved as a result of fluid
unmixing in the runs with NaCl-rich com-
positions. The lower chlorine D- value found
for the experiment with 25 wt % NaCl,
compared to that with 15 wt % NaCl, may
therefore suggest a higher proportion of
this H20-rich fluid in the former experi-
ments. Figure 36 portrays fluid composi-

GEOPHYSICAL LABORATORY

NaCI

Fig. 36. Bulk composition of fluids in the system
C02-H20-NaCl (wt %) for experiments involving
CO2-H2O fluids in which H26 was added as 4-25
wt % NaCI solutions (see caption to Fig. 3 5 for
more details). The tie-lines and phase field bound-
ary are schematic but meant to be consistent with
the observed partitioning relations (see text).

tions for the experiments involving mixed
CO2-H2O fluids with 4, 15 and 25 wt %
NaCI plotted in the C02-H20-NaCl ter-
nary system. Also shown in this figure is a
topology for the two-phase field that is
consistent with the above observations.
Although no previous experimental mea-
surements at high P and Thave been made
with regard to the extent of immiscibility
for C02-H20-NaCl fluids , results at low P-
T conditions (Popp etal., this Report) indi-
cate that the compositional range of the
two-fluid field in this system can be exten-
sive. Experiments employing synthetic
fluid inclusions are now in progress in an
attempt to confirm this interpretation of the
partitioning data.

Partition coefficients for fluorine and
chlorine between apatite and carbonate melt
(obtained at 1050°C, 2.0 GPa) as a function
of total wt % chlorine are shown in Fig. 37.
Fluorine D-values are lower than those
measured for aqueous fluids (i.e., -1.5 vs
>5, respectively), whereas chlorine parti-
tion coefficients (-0.07) are similar to val-
ues measured for experiments with H2O
and HCl-bearing solutions. Partition coef-

Q- 1

.01

Molten Carbonate (Dol80:NaCarb20)
5-1 5 wt % H20

rev

Fluorine

Chlorine

12 3 4 5

Wt % CI

Fig. 37. Apatite/carbonate melt partition coeffi-
cients for fluorine (squares) and chlorine (circles)
as a function of total chlorine concentration (ob-
tained in experiments run at 2.0 GPa and 1050°C).
Data pertain to experiments with 5-15 wt % H2O
in the melt phase.

ficients for both F and CI show little varia-
tion with total CI content (all experiments
had F abundances of -2.0-2.3 wt %). A
reversal experiment that involved a two-
hour preheating of the sample at 1300°C
and 2.0 GPa (thus completely dissolving
the apatite into the melt), then a slow,
isobaric cooling to the final run condition
of 1050°C (and holding there for 48 hours),
yielded identical D-values as in forward
experiments.

Implications for the Fluorine

and Chlorine Content of

Upper-Mantle Fluids

Testimony bearing on the action of flu-
ids in the upper mantle may be found in
certain suites of ultramafic xenoliths that
contain evidence for mineral replacement
by volatile-bearing phases (O'Reilly and

CARNEGIE INSTITUTION

100

meg aery sts
A H20-HCI

A CO 3 Melt

C-bearing (xenoliths)
□ C02-H20-HCI
■ C03 Melt

1 10

Wt%F

Fig. 38. Calculated fluorine and chlorine contents
of high P-T fluids based on the compositions of
mantle-derived apatites. Fluid compositions were
calculated using the partition coefficients mea-
sured in this study for fluids capable of dissolving
appreciable amounts of apatite (i.e., H2O-HCI (±
CO2) fluids or carbonate melt). Apatite composi-
tions from mantle xenolith parageneses were ob-
tained from Wassetal. (1980), Smith etal. (1981)
and Exley and Smith (1982). Carbon-bearing
apatite compositions (xenolith paragenesis) are
from O'Reilly and Griffin (1988) and apatite
megacryst compositions were obtained from
Hervig and Smith (1981) and Irving and Frey
(1984).

Griffin, 1988) or preserve textures indica-
tive of a free vapor (Kovalenko etal., 1987)
or both. Associated with these petrographic
indications for the infiltration of fluids are
elevated concentrations of elements that
are typically present at only low levels in
mantle rocks (e.g., Ba, Cs, Sr and the rare-
earth elements). Analyses of the fluorine
and chlorine content of apatites that occur
as a minor phase in some such fluid-altered
rocks, combined with the results presented
here, may thus provide constraints on the
halogen abundances of some mantle meta-
somatic agents.

Inasmuch as D-values for fluorine and
chlorine were found to depend on fluid
composition, accurate estimates of mantle
fluid halogen contents will, therefore, be
contingent on a judicious choice of parti-

tion coefficients. Based on results of the
apatite solubility experiments of Baker and
WyUie(1990)andAyersandWatson(1991),
the non-silicate fluids capable of transport-
ing the most significant quantities of apa-
tite are molten carbonate and HCl-bearing
aqueous solutions. By applying the D-
values measured for those compositions (at
1050°C, 2.0 GPa) to analyses of apatites
from mantle parageneses (i.e., present as
interstitial grains in xenoliths or as
megacrysts; see caption to Fig. 38 for data
sources), fluorine and chlorine contents of
fluids that may have coexisted with mantle
apatites were calculated and are portrayed
in Fig. 38. Two of the apatites analyzed by
O'Reilly and Griffin (1988) contain car-
bon abundances of -0.3 and -0.9 wt % and
thus may preserve evidence for equilibra-
tion with a carbon-bearing fluid; partition
coefficients for a CO2-H2O-HCI fluid
[X(C02) = 0.3-0.5] were therefore used in
place of the aqueous HC1 values.

Based on the observed P-T dependence
of D-values found in this study, equilibra-
tion of apatite with fluids at lower P-T
conditions than that assumed for these cal-
culations would not effect calculated CI
abundances but calculated F contents might
represent maximum values. Fluid compo-
sitions determined in the above manner
have minimum Cl/F ratios of > 1, F abun-
dances <1 wt % and chlorine concentra-
tions of from ~1 to -20 wt %. The overall
Cl-rich nature of some calculated fluid
compositions may suggest an important
role for chlorine-bearing fluids as agents of
mass transport in the upper-mantle. This
result is in concert with the measurements
of Brenan and Watson (1991), who ob-

GEOPHYSICAL LABORATORY

served significantly greater partitioning of
trace elements into aqueous chloride solu-
tions (coexisting with olivine) relative to
experiments involving pure H2O. In addi-
tion, both Selverstone et al. (1990) and
Philippot and Selverstone (1991) have docu-
mented the occurrence of trace-element
rich brines in Alpine eclogites, and thus
chlorine-rich fluids may also play a role in
trace element mobilization in other high-
pressure environments.

References

Ayers, J. C. and E. B. Watson, Solubility of
apatite, monazite, zircon, and rutile in
supercritical fluids with implications for sub-
duction zone geochemistry, Phil. Trans. R.
Soc. Lond. A, in press, 1991

Baker, M. B. and P. J. Wyllie, High-pressure
solubility of apatite in carbonate-rich melt
(abstr), EOS, 70, 1394, 1990.

Brenan, J. M. and E. B. Watson, Partitioning of
trace elements between olivine and aqueous
fluids at high P-T conditions: Implications for
the effect of fluid composition on trace ele-
ment transport, Earth Planet. Sci. Lett., in
press, 1991.

Candela, P. A., Toward a thermodynamic model
for the halogens in magmatic systems: an
application to melt-vapor-apatite equilibria,
Chem. Geoi, 57, 289-301, 1986.

Exley, R. A. and J. V. Smith, The role of apatite in
mantle enrichment processes and in the
petrogensis of some alkali basalt suites,
Geochim. Cosmochim. Acta, 46, 1375-1384,
1982.

Hervig, R. L. and J. V. Smith, Dolomite-apatite
inclusion in chrome-diopside crystal, Bellsbank
kimberlite, South Africa, Amer. Mineral., 66,
346-349.

Irving, A. J. and F. A. Frey, Trace element abun-
dances in megacrysts and their host basalts:
Constaints on partition coeffcients and
megacryst genesis, Geochim. Cosmochim.
Acta,48, 1201-1221, 1984.

Korzhinskiy, M. A., Apatite solid solution as
indicators of the fugacity of HC1° and HF° in
hydrothermal fluids, Geochem. Int., 18, 44-
60, 1981.
Kovalenko, V. I., I. P. Solovova, I. D. Ryabchikov,
D. A. Ionov, O. A. Bogatikov and V. B.
Naumov, Fluidized C02-sulphide-silicate
media as agents of mantle metasomatism and
megacrysts formation: evidence from a large
druse in a spinel-lhezolite xenolith, Phys. Earth
Planet Int., 45,280-293,1987.
Latil, C. and R. Maury, Contribution a l'etude des
echanges d'ions OH*, CI" et F" et de leur
fixation dans les apatites hydrothermales,5«//.
Soc. Fran. Mineral. Cristall., 100, 246-250,
1977.
O'Reilly , S. Y. and W. L. Griffin, Mantle metaso-
matism beneath western Victoria, Australia:
1. Metasomatic processes in Cr-diopside
lherzolites, Geochim. Cosmochim. Acta, 52,
433-447, 1988.
Philippot, P and J. Selverstone, Trace-element-
rich brines in eclogite veins: implications for
fluid composition and transport during sub-
duction, Contrib. Mineral. Petrol., 106, 417-
430, 1991.
Piccoli, P. M. and P. A. Candela, The mathemati-
cal modeling of the halogen composition of
the mineral apatite during first and second
boiling (abstr), EOS, 72, 312, 1991.
Smith, J. V., J. S. Delaney, R. L. Hervig, and J. B.
Dawson, Storage of F and CI in the upper
mantle: geochemical implications, Lithos, 14,
133-147, 1981.
Tacker, R. C. and J. C. Stormer, A thermodynamic
model for apatite solid solutions, applicable to
high-temperature geologic problems, Amer.
Mineral, 74, 877-888, 1989.
Wass, S . Y. , P. Henderson and C. J. Elliott, Chemi-
cal heterogeneity and metasomatism in the
upper mantle: Evidence from rare earth and
other elements in apatite-rich xenoliths in ba-
saltic rocks from eastern Australia, Phil. Trans.
R. Soc. Lond. A, 297, 333-346, 1980.
W'atson, E. B., Apatite and phosphorous in mantle
source regions: an experimental study of apa-
tite/melt equilibria at pressures to 25 kbar,
Earth Planet. Sci. Lett., 51, 322-335, 1980.
Zhu, C. and D. A. Sverjensky, A set of consistent
thermodynamic properties for fluorapatite,
hydroxylapatite and chlorapatite (abstr), EOS,
72, 145, 1991.

CARNEGIE INSTITUTION

Investigation of Fluid Immiscibility in

the System H2O-NACL-CO2 Using Mass

Spectrometry and Microthermometry

Techniques Applied to

Synthetic Fluid Inclusions

Robert K. Popp? John D. Frantz, and
Thomas C. Hoering

Popp and Frantz (1990) described the
use of synthetic fluid inclusions to define
the limits of fluid miscibility in the system
H20-NaCl-C02. Inclusions were trapped
in quartz prisms contained in fluids with
sodium chloride and carbon dioxide con-
tents up to 1 8 and 50 wt %, respectively, at
temperatures of 500°, 600°, and 700°C and
pressures of 1000, 2000, and 3000 bar. The
presence of two different types of inclu-
sions within a single sample was inter-
preted as evidence of immiscibility for a
particular fluid composition, temperature,
and pressure. Samples that entrapped im-
miscible fluids exhibited both (1) inclu-
sions containing a salt crystal in addition to
a vapor bubble and (2) inclusions without
salt crystals, but with a much larger bubble
than in the first case. The former type
inclusion represents the quenched high-
density (i.e. NaCl-rich, C02-poor) phase,
whereas the latter type represent the
quenched low-density (i.e. C02-rich, NaCl-
poor) phase. Inclusions that formed in ex-
periments in which the fluid was in the
miscible field contained only a single type
that did not contain NaCl crystals and had
proportionally the same ratio of bubble to

Department of Geology, Texas A&M Univer-
sity, College Station, Texas 77843

total inclusion volume. This report de-
scribes the use of mass spectrometric tech-
niques and microthermometric measure-
ments to define more precisely the location
of the solvus separating the one -phase and
two-phase regions in T-P-X space, and to
define the chemical compositions of the
coexisting high-density and low-density
fluid phases.

Mass Spectrometry

A Finnigan 4500 quadrupole mass spec-
trometer was modified (Fig. 39) for the

Fig. 39. Modification of Finnigan 4500 mass
spectrometer. See text for details.

analyses of individual and groups of fluid
inclusions using the general technique of
Barker and Smith (1986), as modified by
Frantz et al. (1989). A crushed sample of
the quartz prism (-20 mg) was placed in a
silica tube (B) surrounded by a resistance
heater (A). The tube, sealed at the outer
end, was inserted into a flight tube (C)
extending through the existing vacuum lock
of the mass spectrometer. The tube is
designed to stretch out the arrival time at
the ion source of the pulses of released
gases resulting from inclusion decrepita-
tions. The flight tube connects with a modi-
fied electron beam cup (G) and focuses
most of the released gas molecules into the
path of the electron-beam where they are

GEOPHYSICAL LABORATORY

390

Temperature, °C

400

Temperature, °C

410
— I

408

418

428

H,0*

_K_»i , — lJ-a

h;5jJXm^^

— COfe

*K~*

jJv.

— C02

.wjs~

' I ■

J I 1_

J I I I 1_

J I I I I i_

21000 21500 22000 22500

Scan Number

i i i i i i i i i i i i i i i i ■'■■

23000 23500 24000 24500 25000

Scan Number

Fig. 40. Mass spectrograms showing H20+ and C02+ intensities for fluid-inclusion decrepitations from
(a) a quartz sample equilibrated at 500°C, 1000 bar with a fluid of composition NaCl7.8-(C02)25.2-
H2C>67.o(wt %) and (b) a quartz sample equilibrated at 500°C, 1000 bar with a fluid of composition
NaCl9.4-H2O80.6-(CO2)i0.0 (wt %). The spectrogram in Fig. 40a demonstrates the existance of two types
of inclusions having different CO2/H2O ratios; the spectrogram in Fig 40b, the existence of one type of
inclusion.

ionized. These ions are then accelerated
and separated by the quadrupole mass filter
(D) and detected by the ion multiplier de-
tector (E). A cryogenic pump (F), chilled
with liquid nitrogen, is used to reduce the
water background at mass 18. Because the
mixtures of low molecular weight gases
analyzed in this study are simple compared
to the organic compounds often analyzed
with the instrument, the selected ion moni-
toring mode of the INCOS 2300 data sys-
tem of the mass spectrometer was used.
Only the selected masses of interest were
measured. A total scan time of 28 millisec-
onds for two masses was utilized.

The mass spectrometer was used to ana-
lyze the ratio of CO2 and H2O in the fluid
inclusions by monitoring channels corre-
sponding to masses 18 (H2O4") and 44
(CO24"). Approximately 45,000 scans were
collected as the samples were heated from
200 to 600 °C at 5 °C per minute. Decrepi-
tation of the inclusions generally occurred
between 300 and 573 °C (the latter being
the approximate temperature of the oc-p
transition for quartz). A spectrogram from

the analysis of inclusions from an experi-
ment at composition NaO7.8-CO225.2~
H2O67.O equilibrated at 500°C and 1000
bar is shown in Fig. 40a. The 2000 scans
shown in the figure correspond to a rise in
the sample temperature from 390° to 410°C
over a time period of approximately 240
seconds. The peaks, which correspond to
the decrepitations of individual and mul-
tiple fluid inclusions, cover a range varying
from 10 to 15 spectrometer scans. Indi-
vidual peaks represent only nanogram
amounts of H2O and CO2. The arrival
times (scan number) and shapes of the CO2
and H2O peaks are quite similar, though the
sensitivity for water tends to be somewhat
less than that of carbon dioxide due to the
tendency of water molecules to absorb on
the surface of the flight tube. Fluid immis-
cibility clearly exists, as evidenced by the
large variability in the area ratios of CO2
and H2O. In the case of fluid of composi-
tion NaCl9.4-CO2i0.0-H2O80.6 equilibrated
at 500°C and 1000 bar (Fig. 40b), a single
fluid phase was present at the experimental

CARNEGIE INSTITUTION

O
0

Mean = 0.406
Std Dev = 0.039

40 60 80 100

Area %, C02

40 60 80 100

Area %, C02

NaCI

H20 10 20 30 40 50

CO-

Fig. 41. Histograms of the frequency of the area % CO2 [areaco2(/areaco2+areaH20)] from mass
spectrograms resulting from decrepitation of fluid inclusions All five samples were equilibrated with
a series of compositions at 500°C, 2000 bar in the single-phase fluid compositional region, as shown
in the ternary diagram in the lower right.

run conditions, because the area ratios of
CO2 and H2O are nearly identical for all the
peaks.

Five samples equilibrated at 500°C and
2000 bar containing less than 5.2 wt %
NaCI with varying CO2 contents were se-
lected for calibration of the spectrometer.

Based on the optical detection of only a
single inclusion type in each sample, at was
concluded that the five samples have trapped
miscible fluids. For all five samples, areas
under corresponding CO2 and H2O peaks
resulting from decrepitations were com-
puted, using the INCOS 2300 software

GEOPHYSICAL LABORATORY

10 20 30 40 50 60

wt % co2

Fig. 42. Calibration curve showing the relation
between area % CO2 measured by mass spectrom-
etry and wt % CO2 of the equilibrated fluid for the
samples shown in Fig. 41. The squares represent
the mean values of the measured area percents
with the brackets indicating one standard devia-
tion.

standard with the Finnigan 4500 spectrom-
eter. Histograms of area % CO2 measured
for the peaks are shown for the five samples
in Fig. 41. Fig. 42 shows the mean values
for all five samples, with their correspond-
ing standard deviations, plotted against the
concentrations of CO2 (in wt %) initially
added to the experimental charges. The
second-order quadratic least squares fit of
these data was then used as the calibration
curve to define the wt% CO2 in inclusions
grown in the two-phase region, for which
C02-contents were unknown.

Eight samples equilibrated at 500°C and
1000 bar demonstrate the results obtained
for inclusions grown in the immiscible
field (Fig. 43). Histograms labelled 1 and 2
are from samples that trapped miscible
fluids so that their mean C02-contents cor-
respond closely to the original CO2 con-
tents of the fluid, denoted by the vertical
dashed lines. The other six histograms,
however, correspond to samples that trapped
immiscible fluids. In histograms 3 through
6, the intervals exhibiting the highest fre-

quency are generally at much lower values
of wt % CO2 than the initial bulk composi-
tion, but some more C02-rich intervals
contain smaller populations. The highly-
populated intervals of low wt % CO2 rep-
resent the concentration of cabon dioxide
in the sodium chloride-rich, high-density
fluid phase. In the case of histogram 7, for
which the original bulk fluid composition
lies extremely close to the carbon dioxide-
rich limb of the solvus, the interval of
highest frequency lies in the region of the
original bulk composition, with less popu-
lated intervals lying at lower CO2 concen-
trations. In this case, the highly populated
intervals correspond to the concentration
of carbon dioxide in the C02-rich, low-
density fluid phase. Histogram 8
demonstates a case in which highly popu-
lated intervals extend from the original
bulk composition down to those represent-
ing quite low concentrations of CO2. Be-
cause the two immiscible phases are likely
to be intimately intermixed rather than sepa-
rated into a single high-density and a single
low-density phase within the capsule
(Ramboz et al., 1982; Zhang and Frantz,
1989), many inclusions trap varying pro-
portions of the two fluids; thus, a range of
inclusions with properties intermediate
between the two end members is com-
monly observed in a single sample from the
immiscible field. The less-populated inter-
mediate intervals result from decrepita-
tions of inclusions containing mixtures of
the two fluids and from simultaneous de-
crepitations of groups of vapor-rich and
liquid-rich inclusions.

The use of mass spectrometry measure-
ments resulted in an expanded two-phase

CARNEGIE INSTITUTION

0 20 40 60 80 100

Wt%C02

NaCI
/

l^£v

0 20 40 60 80 100

Wt%C02

H2Q 10 20 30 40 50 C02

Fig. 43. Histograms of the frequency of the values of wt % CO2 from mass spectrograms resulting from
decrepitation of fluid inclusions grown at 500°C, 1000 bar. The values of area % CO2 were converted
to wt% CO2 using the calibration curve shown in Fig. 42.

NaCI +
Solution

20 30 40 50 60

Wt % NaCI

Fig. 44. Solubility diagram (in wt %) of NaCI in
NaCl-H20 solutions as a function of temperature
(data from Linke, 1958, 1965).

field relative to that obtained from only the
optical identification of the two inclusion
types. That is, the mass spectrometry tech-
nique has the increased precision neces-

sary to identify inclusions of more than one
composition in some samples where only
one inclusion type has been identified op-
tically. In addition, the technique permits
the CO2 content of the high-density fluid to
be determined. The concentration of CO2
in the high-density fluid is taken to be that
of the highest frequency intervals at rela-
tively low weight % CO2 (Fig. 43).

Microthermometry

In order to determine the positions of
tie-lines connecting the compositions of
coexisting immiscible fluids, the concen-
tration of NaCI in the high-density fluid
inclusions was estimated from

GEOPHYSICAL LABORATORY

500 °C
2000 bar

500 °C
1000 bar

H20 10 20 30 40 50

£Q H20 10 20 30 40 50 C02

NaCI

600 °C

1000 bar __ f

H~0 10 20 30 40 50

700 °C

1 000 bar 60

CO H2° 10 20 30 40 50

CO.

Fig. 45. Compositions of coexisting immiscible fluids. Each tie line represents
the compositions of the coexisting fluid as determined for the sample denoted
by the filled circle through which they pass. See text for further details.

microthermometry measurements made
using a heating stage supplied by Fluid
Inc., Denver, Colorado, In the phase dia-
gram shown in Fig. 44, the NaCI content of
a given bulk composition is known if the
temperature of the univariant boundary
between the "Solution" field and the "NaCI
+ Solution" field is known (i.e., if the
temperature of NaCI melting is known).
The presence of CO2 in the fluid inclusions
might affect the temperatures obtained from
the phase relations in the C02-free system.
However, the results of the mass spectrom-

etry measurements described above sug-
gest that the bulk concentration of carbon
dioxide in the high-density inclusions is
relatively low. In addition, most of that
carbon dioxide is contained in the vapor
phase (i.e., the bubble) at the temperatures
of NaCI melting. Therefore, only very
small CO2 concentrations must be con-
tained in the liquid phase, and thus the
effect of CO2 on the temperature of NaCl-
melting is considered insignificant.

The melting temperatures of NaCI crys-
tals in the high-density inclusions were

CARNEGIE INSTITUTION

determined for selected samples equili-
brated at 500°C, 1000 and 2000 bar; 600°C,
1000 bar; and 700°C, 1000 bar. Arelatively
large variation in the melting temperature
was observed in each individual sample,
with the largest frequency occurring at the
highest temperatures. The lower-tempera-
ture measurements are obtained from in-
clusions that trapped mixtures of the two
immiscible fluids, and therefore contain
lower NaCl concentrations than inclusions
that trapped only the high-density phase.
With knowledge of NaCl contents, ob-
tained from the highest NaCl melting tem-
perature measured in a given sample, and
knowledge of the CO2/H2O ratios mea-
sured by mass spectrometry, the chemical
composition of the high-density fluid phase
is completely defined. Tie lines were lo-
cated between the coexisting fluid phases
at 500°Cand2000bar,500°Cand lOOObar,
600°C and 1000 bar, and 700°C and 1000
bar (Fig. 45). To construct the tie lines, a
straight line was drawn from the composi-
tion of the high-density fluid through the
bulk composition to the carbon dioxide-
rich limb of the solvus. At a given tempera-
ture and pressure, the slopes of the tie lines
systematically steepen as the NaCl-H20
binary is approached. With increasing tem-
perature, the tie-lines become increasingly
steeper, as a result of greater partitioning of
sodium chloride and reduced partitioning
of carbon dioxide between the two immis-
cible phases.

Summary
The use of mass spectrometry in

conjunction with the more routine optical
and microthermometric techniques pro-
vides an improved basis on which to sepa-
rate fluid inclusions that trapped one -phase
fluids from those that trapped two-phase
fluids. In addition, the technique provides
a relatively straightforward method to de-
fine tie lines between the compositions of
the coexisting high-density and low-den-
sity fluids. The techniques described here
should be applicable to other systems of
petrologic interest, such as those contain-
ing the volatile species CH4, N2, and HC1.

References

Frantz, J.D., Y. Zhang, D. D. Hickmott, and T. C.
Hoering, Hydrothermal reactions involving
equilibrium between minerals and mixed
volatiles. 1 . Techniques for experimentally load-
ing and analyzing gases and their application to
synthetic fluid inclusions, Chemical Geol., 76,
57-70, 1989.

Barker, C. and M. P. Smith, Mass spectrometric
determination of gas in individual fluid inclu-
sions in natural minerals. Anal. Chem., 58,
1330-1333, 1986.

Linke. W.F., Solubilities of Inorganic and Metal-
Organic Compounds, 1. Van Norstrand,
Princeton, N.J., 4th ed., 1958.

Linke. W.F., Solubilities of Inorganic and Metal-
Organic Compounds, 2. Am . Chem . Soc . , Wash-
ington D.C, 4th ed., 1965.

Popp, R.K. and J.D. Frantz, Fluid immiscibility in
the system H20-NaCl-C02 as determined from
synthetic fluid inclusions, Annu. Rep. Director
Geophys. Lab., Carnegie Instn. Washington,
1989-1990,43-47, 1990.

Ramboz, C, M. Pichavant, and A. Weisbrod,
Fluid immiscibility in natural processes: Use
and misuse of fluid inclusion data, II. Interpre-
tation of fluid inclusion data in terms of immis-
cibility, Chemical Geol, 37, 29-48r 1982.

Zhang, Y. and J. D. Frantz, Experimental determi-
nation of the compositional limits of immisci-
bility in the system CaCl2-H20-NaCl at high
temperatures and pressures using synthetic fluid
inclusions, Chemical Geol., 74,289-308, 1989.

GEOPHYSICAL LABORATORY

The Akermanite-Gehlenite-Sodium

Melilite Join at 950°C and 5 kbar in the

Presence of CO2 + H2O

H.G. Huckenholz, H.S. Yoder, Jr., T.
Kunzmann* and W.Seiberl*

Melilites are primarily solid solutions
between akermanite (Ca2MgSi20y), so-
dium melilite (NaCaAlSi207), and
gehlenite (Ca2Al2Si07). High-grade meta-
morphism of impure limestones and dolo-
mites favors the crystallization of members
of the akermanite - gehlenite solid solution
series, and the sodium melilite component
is generally low. During the decarbonation
process, CO2 and H2O play an active role
and greatly influence the metamorphic as-
semblages. On the other hand, igneous
rocks, usually highly undersaturated and
alkalic, are generally enriched in the so-
dium melilite component and are close to
the akermanite - sodium melilite join near
ak60ge 1 osm30 0^ Goresy and Yoder, 1 974).
A C02-rich fluid is probably involved in
the upper mantle melting process, and is
thought by some to be responsible for the
formation of melilitite magma as well as
nephelinite and kimberlite magmas.

Experimental data on melilite-C02-H20
are available for the endmember akermanite
(Yoder, 1968, 1973, 1975; Huckenholz et
al., this Report), gehlenite (Huckenholz
and Yoder, 1974; Huckenholz, 1977;
Hoschek, 1974) and for the akermanite -
gehlenite solid solution series (Huckenholz

Mineralogisch-Petrographisches Institut,
Ludwig-Maximilians Universitat, D-8000
Miinchen 2, Germany.

et al., 1990). In order to elucidate the
crystallization behavior of ternary melilites
and their relationship to adjoining phases,
an experimental study of the isothermal,
isobaric section at 950°C and 5 kbar of the
akermanite - gehlenite - sodium melilites
was conducted in the presence of H2O and
CO2 in both Washington and Munich.

Experimental procedure

Thirty crystalline melilite compositions,
prepared by J.F. Schairer, from the
akermanite - sodium melilite and the
akermanite - gehlenite - sodium melilite
joins (Schairer and Yoder, 1964; Schairer et
al., 1967) were used in the experiments.
Ten crystalline melilite compositions were
also available from the akermanite -
gehlenite join (Huckenholz et al., 1990)
and (pure) sodium melilite from Yoder 's
(1973) experimental study on that
endmember composition. Water was added
in excess to the samples for the melilite-
H2O experiments, whereas oxalic acid
dihydrate (H2C204»2H20) served as a
source of CO2 + H2O. Oxalic acid dihydrate
produces an initial X(C02) [CO2/
(H2O+CO2)] of 0.5. Higher X(C02) were
obtained by adding calcite + oxalic acid
dihydrate to the sample, and lower X(C02)
were generated by a mixture of oxalic acid
dihydrate + water. The weight ratio of
oxalic acid dihydrate to sample was about
1:2-3. When carbonation of sample took
place, that is, formation of calcite and sea-
polite, a final X(C02) of about 0.30 ± 0.03

CARNEGIE INSTITUTION

Table 1 1. Liquid compositions

Liquid +

H20

Liquid +

CO2 + H2O

Sm 100-1*

Sm 100-2

Sm8

SmlOO

Sm5

SmlO

Sm67

Si02

43.40

42.57

44.44

48.01

48.60

50.03

48.80

AI2O3

22.98

22.08

(23.00)

21.30

23.16

20.27

(21.70)

Cat)

0.03

0.50

0.06

0.51

0.54

0.50

15.59

15.65

(15.60)

9.30

8.55

9.18

(9.10)

Na20

8.18

9.25

8.40

11.49

11.11

10.02

10.53

totals

90.18

89.58

(91.94)

90.16

91.93

90.04

(90.63)

mel

48.4

53.1

47.9

39.6

23.9

29.1

27.4

(ak)

6.6

0.4

3.6

3.8

3.4

(ge)

27.2

19.5

24.8

7.4

13.4

10.7

12.2

(Sm)

21.2

33.6

16.5

31.8

6.9

14.6

11.8

23.2

36.0

26.8

39.2

55.5

31.3

44.4

28.4

10.9

25.3

21.2

20.6

39.6

28.2

X(CQ2)V

0.0

0.72

0.63

0.54

0.40

Compositions of glasses from 950°C and 5 kbar experiments. SmlOO- 1, sodium melilite
(starting from glass); Sm 100-2, sodium melilite (starting from crystalline sodium melilite);
Sm8, akermanitel0-gehlenite20- sodium melilite70; Sm5, akermanite5-sodium melilite95;
SmlO, akermanite 10- sodium melilite90; Sm67} akermanite67- sodium melilite33. Molecu-
lar proportions are (ak) Ca2MgSi207; (ge), Ca2Al2SiC>7; (Sm), NaCaAlSi207; jd,
NaAlSi206; ab, NaAlSi308- ** Numbers in parentheses are estimates.

Table 12. Compositions of crystalline materials

Liquid + H20

Liquid

+ CO2 + H2C

SmlOO

Sm8

SmlOO

Sm5

Sm64

Sm67

AklOO

mel

1.002

1.716

0.004

0.019

0.001

0.000

0.002

1.059

0.001

0.015

0.000

0.001

0.000

0.001

0.250

0.000

0.002

0.003

0.002

0.995

1.547

0.992

0.961

0.996

0.994

0.997

0.001

0.432

0.003

0.001
Liquid + C02 +

0.001
H20

0.001

Sm5

SmlO

Sm20

Sm67

Sm64

Sm67

AklOO

cpx

0.991

1.001

0.998

1.918

1.925

1.996

0.003

0.007

0.002

0.226

0.190

0.003

0.008

0.024

0.011

0.014

0.848

0.876

0.997

0.996

0.972

0.984

0.979

0.989

1.002

0.002

0.005

0.001

0.029

0.020

0.002

Compositions of wollastonite, melilite, calcite, and clinopyroxene from 950°C and 5 kbar
experiments. Compositions in cations per formula units. Bulk compositions as in Table 1 1 ;
others are: Sm64, akermanite64-sodium melilite36; Sm20, akermanite20- sodium meliliteso;
aklOO, akermanite 100.

GEOPHYSICAL LABORATORY

950"C-5 kbar

Sm
NaCaAISip7

L+wo+v

Ge
Ca2AI2Si07

Sm
NaCaAISi^

L+cc+V

+mel+cc+wo+V

Ge
CaAI^SJaG;

L+mel+cc+cpx+wo+V

Ca^AgS\p7

Fig. 46. Ak-Ge-Sm-H20 isothermal-isobaric sec-
tion at 950°C and 5 kbar. Line A -Bis the limit of
solid solution for ternary melilites found in natural
rocks (El Goresy and Yoder, 1974). Abbrevia-
tions are: V, fluid; L., liquid; mel, melilite; wo,
wollastonite; gar, garnet; and cpx, clinopyroxene.
Symbols refer to the phase assemblages as la-
belled in the plane. Circle with dot is the melilite
composition analyzed.

in the subsolidus assemblages resulted.

In hypersolidus assemblages, H2O par-
titions between fluid and liquid. For the
liquid, an X(C02) between 0.05 and 0. 1 is
assumed. This assumption is derived from
mass balance calculations conducted on a
glass quenched from the sodium melilite
endmember composition in which calcite
is the only liquidus phase, and from CO2 +
H2O solubility in albite melts (Kadik and
Eggler, 1974). The CO2 partitions pre-
dominantly between fluid and calcite +
scapolite. The "final" X(C02) that results
from X(C02)(fluid) + X(C02)(liquid) can
be calculated, and is about 0.3 in most of
the hypersolidus experiments.

Experiments were carried out by means
of two internally heated gas-media appara-
tus in Washington and in Munich. The run
duration was 24 hours in each case.
Reversibility of experiments has not as yet

Fig. 47. Ak-Ge-Sm-H20-CC>2 isothermal-iso-
baric section at 950°C and 5 kbar at an X (CO2) of
about 0.3. Abbreviations as Fig. 46; other is scap,
scapolite. Symbols refer to the phase assemblages
as labelled on the diagram. Square with dot is the
phase assemblage at the beginning of melting.
Line A - B is the position of the X(CC>2) versus
composition plot of data for Fig. 48B.

been demonstrated for this preliminary re-
port. The run products were examined by
optical methods, x-ray powder diffraction,
and for 10 compositions, by microprobe
analyses. Composition of glasses, calcite,
wollastonite, clinopyroxene, and melilite
are given in Tables 11 and 12.

Experimental Results

The experimental data obtained are
presented in two isothermal-isobaric (950C
- 5 kbar) sections of the ak (akermanite)-ge
(gehlenite)-Sm (sodium melilite) join with
H2O and with CO2 + H2O as fluids (Figs.
46 and 47). The two ternary joins are each
part of the cc (calcite)-di (diopside)-CaTs
(Ca-Tschermak's component)-jd (jadeite)
tetrahedron, which in turn is a four-compo-
nent volume of the Na20-CaO-MgO-

CARNEGIE INSTITUTION

950"C-5 kbar

scap (mei) 4cc*cpx tsp.V

■ T T -1 1 1-

| 950'C-5 kbar |

cpx.cc
+wo+V

■

scap+cc*
sp<?)+c<x+V

scap+cc+cpx*sp(?)+V

0*5

mel+cc ^

\ mal+cpx+cc+V

\ ■

■

\ .

md+V

■

10 20

40 50

90 A

/Sm15)

mol%

(Ak85/Sm15)

Al203-Si02-C02-H20 system. Phases
determined are fluid (V), liquid (L) calcite
(cc,CaC03), wollastonite (wo, CaSi03),
clinopyroxene [cpx,

(Ca,Na)(Mg,Al)(Si,Al)206], garnet [gar,
(Ca,Mg)3Al2Si30i2L scapolite [scap,
Ca3(Ca,Na)iAl5-6Si6-7024(C03)] and
melilite [mel, (Ca,Na)2(Mg,Al)
(Si,Al)207]. Compositions of calcite,
clinopyroxene, and melilite plot in the cc-
di-CaTs-jd system, whereas wollastonite,
garnet, scapolite, and the liquid lie outside.
Chemographical correlation of spatial phase
assemblages from the multicomponent sys-
tem Na20-CaO-MgO-Al203-Si02-C02-
H2O, however, cannot always be appropri-
ately illustrated.

The isothermal-isobaric section of Fig.
46 displays the phase assemblages encoun-
tered in experiments with H2O as the fluid
phase. The assemblages are outlined by
dashed curves that are thought to be inter-
faces of multiphase volumes intersecting
the pseudo-ternary plane. The maximum
extent of melilite solid solution, equivalent

V?:

950-C-5kbar

, V+cryilaU+LxMd

cpx

scap

Lcfjd.c/yilals ± cak.«e -'

Y ■ ' 1H9 *ec «cpx jx .3 -

me)(1) .v

o.ot

SO
mol%

Fig. 48. Melilite composition versus AXCO2) .A.
Melilite composition versus X(C02) plot of data
along the join akermanite - gehlenite. Abbrevia-
tions as in Figs. 46 and 47, others are mei, meionite;
cor, corundum; sp, spinel. Symbols refer to the
phase assemblages as labeled on the diagram. B.
Melilite composition versus X(CC>2) plot of data
along line A (Ak85/Sml5) -B (Ge 85/Sm 15) of
Fig. 47. Melilite composition from above and
below that line are projected. Abbreviations as
Figs. 46 and 47; symbols refer to the phase assem-
blages labelled. C. Melilite composition versus
X(C02) plot of data along the join akermanite -
sodium melilite. Abbreviations as in Figs. 46 and
47. Mel (2) lies off the plane in the gehlenite - rich
portion of Fig. 47. Vertical solid lines with dots
are tie lines between fluid and liquid in composi-
tions along akermanite - sodium melilite. Sym-
bols in the subsolidus portion refer to phase assem-
blages as labeled on the diagram.

GEOPHYSICAL LABORATORY

to the H20-saturated solidus, is bound by
akermanite58 sodium melilite42;
akermanite26gehlenite3 1 sodium melilite43
(microprobe data, Table 12; circle with
dot); and by gehleniteyo sodium melilite3o
compositions. The H20-saturated melilite
solidus is close to the extent of melilite
solid solution found for natural melilites by
El Goresy and Yoder (1974). Solid phases
in the hypersolidus portion of the plane
coexist with a water-saturated liquid.
Glasses quenched from sodium melilite ioo
as well as akermaniteio gehlenite20 so-
dium melilite70 were analyzed by electron
microprobe. Their compositions (Table
11) contain the melilite components (= ak +
ge + sm) on the order of about 50 mol % but
also jadeite (NaAlSi206) and albite
(NaAlSi308). The total of constituents
analyzed of the glasses run about 89-91%;
the remainder is believed to be H2O dis-
solved in the melt. The liquidus-phase
wollastonite was also analyzed (Table 12),
and is very close to wollastonite with minor
solid solution, if any. No nepheline was
detected in the run products, neither by
optical, x-ray analysis, nor microprobe in-
spection.

The isothermal-isobaric melilite sec-
tion at 950°C and 5 kbar with CO2 + H2O
as a fluid is depicted in Fig. 48. For simplic-
ity, the ternary plane is averaged forX(C02)
of about 0.3. The bounding melilite solid
solution on the CO2 + H20-saturated solidus
was assumed to be akermanite20-
gehlenite50-sodium melilite30. The appar-
ent boundaries of the phase assemblages
mapped in the plane tend to converge about
this composition.

Further indication of this special melilite

composition can be deduced from the iso-
thermal-isobaric, X(C02) vs. melilite com-
position plot. Figure 48A displays such a
relationship for akermanite-gehlenite. The
solid solution is bound by decomposition
of akermanite + CO2 to diopside + calcite
at 0.12±0.02 and by gehlenite + CO2 to
meionite + calcite + corundum at X(C02)
0.30+0.02 (Huckenholz et al., 1990;
Huckenholz and Seiberl, 1990), respec-
tively. The breakdown assemblages of
melilite + scapolite -1- calcite + corundum +
spinel (?) in the gehlenite-rich portion and
that of melilite + clinopyroxene + calcite in
the akermanite-rich portion indicate a maxi-
mum melilite solid solution of about
akermanite 1 5 -gehlenite85, which appears
to decompose at about X(C02) = 0.33 to
scapolite + calcite + corundum + spinel (?).

A C02-dependent compositional maxi-
mum is also indicated when the relations
are considered along line A -B on the
akermanite-gehlenite-sodium melilite plane
parallel to akermanite-gehlenite at about
sodium melilite 15. In the section displayed
in Fig. 48B, the melilite -1- CO2 stability has
increased to about 0.45 AXCO2) . No stable
melilite was found in the subsolidus assem-
blage of scapolite + calcite + clinopyroxene
+ spinel (?) at 0.5, and above the scapolite
is a solid solution between meionite
[Ca4Al6Si6024(C03)] and the (theoreti-
cal) carbonate-marialite endmember
[Na3CaAl3Si9024(C03)]. At 950°C and 5
kbar its composition in (scapolite + calcite)
- bearing assemblages is restricted to an
equivalent an-contentof 0.75 (Huckenholz
and Seiberl, 1990).

Phase relations along the join
akermanite-sodium melilite as a function

CARNEGIE INSTITUTION

of X(C02) are shown in Fig.48C. The
extent of solid solution toward sodium
melilite is limited to akermanite70-sodium
melilite30 and an X(C02) of 0. 1 2, which is
at the breakdown of akermanite + CO2 to
diopside + calcite. Melilites labelled mel
(2) occur within the phase assemblages
generated by the breakdown of akermanite
+ CO2, but do not lie in the pseudo-binary
akermanite-sodium melilite plane. Their
composition is located in the gehlenite-rich
portion of the akermanite-gehlenite-sodium
melilite join as shown in Fig. 47, with
X(C02) up to about 0.45. From the bound-
ing melilite solid solution of the fluid-
saturated solidus toward the sodium melilite
endmember, hypersolidus phase assem-
blages of V+mel+cc+cpx+scap+L,
V+mel+cc+cpx+wo+L, V+mel+cc+wo+L,
V+mel+cc+L and V+cc+L are traversed by
the pseudo-binary join. Liquid with an
assumed X(C02)(L) of 0.07 coexists with
fluid having X(C02)(V) of 0.45 in the
akermanite-rich portion of the join. To-
ward the sodium melilite endmember com-
position, the amount of the liquid increases
thereby consuming increasing amounts of
H2O, which in turn results in raising the
X(C02) of the fluid.

Quenched glasses from sodium
melilite 100, akermanite5 -sodium melilite95,
akermaniteio-sodium melilite^, and
akermanite67-gehlenite33 bulk composition
have been analyzed by microprobe (Table
11). They exhibit (calculated molecular) jd
(NaAlSi2C>6) and ab (NaAlSi308) compo-
nents in addition to a large amount of
ternary melilite (ak+ge+sm). In contrast,
glasses quenched in the presence of a (pure)

H2O fluid are lower in the jd and ab com-
ponents but contain a ternary mel-compo-
nent > 47%. Analyzed calcite (Table 12)
contains minor amounts of MgO and Na20;
wollastonite displays minor solid solution
toward the sodium melilite (<1%) and
akermanite (2-4%) components.
Clinopyroxenes were analyzed from the
akermanite64-sodium melilite36,
akermanite67-sodium melilite33, and
akermaniteioo bulk compositions. Ex-
pressed as endmembers, they reduce to
diopside87-CaTsio jadeite3, diopside89-
CaTs9 jadeite2, and diopside99.7-CaTs<o.i
jadeite<o.2, respectively.

Reference

El Goresy, A., and H. S. Yoder, Jr., Natural and
synthetic melilite compositions. Carnegie
Instn. Washington, Year Book, 73, 359-371,
1974.

Hoschek, G., Gehlenite stability in the system
CaO-Al203-Si02-H20-C02- Contr. Min.
Petr., 47, 245-254, 1974.

Huckenholz, H. G., Gehlenite stability relations in
the join Ca2Al2SiC>7 - H2O up to 10 kbar.
NJb. Miner, Abh., 130, 169-186, 1977.

Huckenholz, H. G., and H. S. Yoder, Jr., The
gehlenite-H20 and \vollastonite-H2O systems.
Carnegie Instn. Washington, Year Book, 73,
440-443, 1974.

Huckenholz, H.G., A. Wassermann, and K. T.
Fehr, Stability and phase relations of gehlenite-
akermanite solid solutions in the presence of a
H20-C02-fluid. International Symposium of
Experimental Mineralogy, Petrology and Geo-
chemistry, Edinburgh, UK, p. 17, terra ab-
stracts, 2, 1990.

Huckenholz, H. G., and W. Seiberl, Stability and
phase relations of carbonate scapolite solid
solutions under the PT-regime of the deeper
crust. Third International Symposium of Ex-
perimental Mineralogy, Petrology and Geo-
chemistry, Edinburgh, UK,. P. 17; terra ab-
stracts 2 1990.

Kadik, A. A., and D. H. Eggler,Melt-vapor rela-
tions on the join NaAlSi308-H20-C02:
Carnegie Instn. Washington, Year Book, 74,
479-484, 1974.

GEOPHYSICAL LABORATORY

Schairer, J. R, and H. S. Yoder, Jr., The join
akermanite (Ca2MgSi207) - soda melilite
(NaCaAlSi207) . Carnegie Instn. Washing-
ton, Year Book, 63, 89-90, 1964.

Schairer, J.F., H. S. Yoder, Jr., and C. E. Tilley,
The high-termperature behavior of synthetic
melilites in the join gehlenite-soda melilite-
akermanite, Carnegie Instn. Washington, Year
Book, 65, 217-226 1967.

Yoder, H. S., Jr., Akermanite and related melilite-
bearing assemblages. Carnegie Instn. Wash-
ington, Year Book, 66, p47 1-477, 1968.

Yoder, H. S. Jr., Melilite stability andparagenesis.
Fortschr. Mineral, v.50, 140-173, 1973.

Yoder, H.S., Jr., Relationship of melilite- bearing
rocks to kimberlite: a preliminary report on the
system akermanite-C02. Proc. Internat.
kimberlite Conf., Cape Town. Phys. Chem.
Earth, 9, 883-894, 1975.

Merwinite Stability and

High-Temperature Phase Relations

in the Presence of CO2 + H2O.

H. G. Huckenholz, H. S. Yoder, Jr., and
W. Seiberl

monticellite +melilite-bearing assemblages
in high-temperature calc-silicate rocks oc-
curring as inclusions in pyroxenites from
the critical zone of the eastern Bushveld
Complex.

Phase equilibria studies on the join
CaMgSi206-CaC03-C02 of the CaO-
MgO-Si02-C02 system restrict the
merwinite + C02 stability to high-tem-
perature but low-pressure conditions. The
reaction akermanite + calcite <=> merwinite
+ CO2 (step 1 1 of the decarbonation series
of Bowen, 1940) was studied by
Shmulovich (1969) and Walter (1963a,b,
1965) at low pressure but high tempera-
ture. Merwinite + CO2 crystallizes from
akermanite + calcite assemblages at tem-
peratures of < 1065°C and pressures of <
0.5 kbar. Merwinite + CO2, however, did
not crystallize from diopside + 2 calcite
assemblages between 950°C and 975°C at
1 kbar (Yoder, 1975).

Merwinite [Ca3Mg(Si04)2] was discov-
ered and named by Larsen and Foshag
(1921) from high-grade metamorphosed
carbonaceous rocks at Crestmore near Riv-
erside, California. At Crestmore (Burnham,
1959; Walter, 1965) and other localities
(e.g., Scawt Hill, Northern Ireland, Tilley,
1929; Ardnamurchan, western Scotland,
Agrell, 1965; Christmas Mountains, Big
Bend region, Texas, Joesten, 1974),
merwinite is mainly associated with cal-
cite, spurrite, monticellite, melilite, and
also with larnite. Recently, Wallmach et al.
(1989) described merwinite from

Mineralogisch-Petrographisches Institut,
Ludwig-Maximilians Universitat, D-8000
Munchen 2, Germany.

Experimental Methods

In order to elucidate the stability of
merwinite in the presence of CO2 + H2O at
pressures > 0.5 kbar, an experimental study
of merwinite and its high-temperature phase
relations was conducted. Stability and phase
relations of merwinite with akermanite,
diopside, calcite, liquid, and fluid were
studied for pure CO2 in the 1-10 kbar
pressure range at temperatures between
900° and 1450°C. In addition, isobaric
temperature versus CO2 + H2O relations
were investigated at 1 and 3 kbar and at
temperatures between 700° and 1 200°C.

Mixtures of crystalline materials were
used in all cases for the experiments. They
consisted of:

CARNEGIE INSTITUTION

03
_Q
J*:

13
CO
CO

12
11

ioh

9
8

7
6
5
4
3
2
1

800

0 o /•/•/• •/«>

900

1000 1100 1200

Temperature, °C

1300

1400

Fig. 49. Pressure-temperature diagram for merwinite + CO2, akermanite + calcite, and diopside + calcite
compositions. Heavy lines are univariant reaction curves; light lines are restricted reaction curves; short
dashed curve is the compositional singularity for di + 2 cc = L; I\, I2, invariant points; Si, 52 singular
points. Abbreviations for phases are: mer, merwinite (Ca3MgSi20g); ak, akermanite (Ca2MgSi2(>7);
di, diopside (CaMgSi2C>6); cc, calcite (CaCC>3); L, Liquid. Symbols: Solid triangles (1) diopside +
calcite = akermanite + CO2 (Shmulovich, 1969; Walter, 1963); solid triangles (2) akermanite + calcite
= merwinite + CO2 (Shmulovich, 1969; Walter, 1963); open hexagons, di + cc from mixture B, C, and
Y (for composition see text); open squares, akermanite + CO2 from mixture B, C, and Y as well as
Yoder's data (1975); open diamonds, merwinite + CO2 from mixture B, C, and Y; solid diamonds, mer
+ L + CO2 from mixtures B, C, and Y; solid squares, ak -1- L and ak + L + CO2 from mixture B, C, and
Y; solid pentagon and solid circle are the 2 kbar run data on di + ak + CO2 = L (Yoder, 1975); solid
hexagons, di + L and Di + L + CO2 from mixture C, Y, and Yoder's 1975 run data; open circles, Liquid
on diopside + cc -1- CO2 compositions, mixture B, C, and Yoder's (1975) run data.

(1). merwinite (crystallized at 1200°C, 1
atm), mixture A;

(2). akermanite (crystallized at 1050°C at 1
atm) + calcite (Baker Chemical Com-
pany, grade C.R.) equivalent (on a mo-
lar basis) to merwinite + CO2; mixture
B;

(3). natural diopside (Twin Lakes, Califor-
nia; Smith, 1 966) + 2 calcite equivalent

to akermanite + calcite + CO2 or
merwinite -1- 2 CO2, mixture C; and

(4). natural 2 wollastonite (Willsboro, N. Y.)
+ natural dolomite (Thornwood, N. Y.)
equivalent to akermanite + calcite +
CO2, or diopside + 2 calcite, or
merwinite + 2 CO2, mixture Y.
Other experiments were carried out on

natural rock inclusions from the upper zone

GEOPHYSICAL LABORATORY

of the eastern Bushveld Complex (locality
Luipershoek, Joubert, 1976). Rocks con-
taining akermanite + diopside (±
monticellite), Ji, and monticellite +
akermanite (± diopside), J6, were collected
by H. G. Huckenholz (October, 1990). The
synthetic assemblages merwinite +
monticellite + akermanite60-gehlenite40
(mixture E) and akermanite + monticellite
+ calcite (mixture D) were also studied at 1
and 3 kbar in the presence of CO2 + H2O.
Experimental data were obtained by means
of internally-heated, gas-media apparatus
in both Washington, D. C, and Munich.
Reversibility of the experiments was en-
sured by the direction of reaction of the
different crystalline mixtures listed above.

Merwinite Relations with CO2

Stability of merwinite -1- CO2 and
merwinite phase relations with akermanite,
diopside, calcite, liquid, and CO2 are dis-
played in the temperature versus pressure
plot in Fig. 49. With the data of Shmulovich
(1969) at 0.5 kbar and below (see also
Walter, 1963a,b), the reaction akermanite +
calcite = merwinite + CO2 increases in
temperature from 1015°C at 0.5 kbar up to
1 1 82°C at 1 .3 ± 0.2 kbar, and results in the
invariant assemblage (h) of merwinite +
akermanite + diopside + calcite + liquid +
CO2 in equilibrium. Four other univariant
curves,

[ak] merwinite + calcite + CO2 <=> liquid,

[mer] akermanite + calcite + CO2
<=> liquid,

[cc] merwinite + CO2

<=> akermanite -1- liquid, and

[CO2] liquid + merwinite

<=> akermanite + calcite,

meet at that invariant point I2. The reaction
[ak] was bracketed between 1150°C and
1200°C at P = 1 kbar. Its position at about
1180°C is fixed due to the (positive) slope
of reaction [mer] and by the run at 1 200°C
and 3 kbar in particular, which is just slightly
above the solidus with a phase assemblage
of akermanite + calcite + liquid. At 2 kbar
and temperatures between 1200° and
1450°C, merwinite + CO2 does not crystal-
lize from akermanite + calcite nor from
diopside + 2 calcite mixtures. Thus, the
curves for the two univariant reactions of
[cc] and [CO2] must pass between 1.3 and
2 kbar. At reaction [cc], merwinite + CO2
tie lines are interrupted by those of
akermanite + liquid but with merwinite
remaining in the C02-absent region of the
merwinite + calcite + liquid and merwinite
+ akermanite + liquid assemblages.

Because of chemographic constraints,
reaction [CO2] must occur on the high-
pressure side of reaction [cc] and between
1.3 and 2 kbar as well. In that pressure
range, the melting curve of calcite (Yoder,
1 973) intersects the reaction [CO2] at about
1350°C. Calcite in the assemblage will
melt and the restricted assemblage of
merwinite + akermanite + liquid evolves
from S\. At higher temperatures, the diop-
side + akermanite + CO2 solidus and the
diopside + CO2 solidus appear in the diop-
side + akermanite + CO2 portion of the
diopside -1- calcite + CO2 system. The
temperature of the akermanite + diopside +
CO2 solidus can be deduced from the run
data of Yoder (1975) on diopside + calcite
compositions obtained at 2 kbar.
Akermanite + diopside + CO2 will melt
slightly above 1400°C, and diopside + CO2

CARNEGIE INSTITUTION

(Rosenhauer and Eggler, 1975) at 1415°C
as well.

In the temperature and pressure range
studied, merwinite and diopside do not
coexist in the presence of C02. They are
separated by akermanite + calcite, or by
akermanite + CO2 tie lines. Below 900°C
there is only a very narrow CO2 pressure
range of about 200 to 300 bar where
akermanite + CO2 is formed from diopside
+ calcite. With increasing temperature,
akermanite + CO2 is stable up to about 5
kbar. The reaction diopside + calcite =
akermanite + CO2 (step 8 of Bowen's de-
carbonation series) was studied at low pres-
sures by Walter (1963) and up to 6 kbar by
Yoder (1973), who exclusively used a crys-
talline mixture of (natural) diopside -1- 1
calcite. The slope of the reaction curve of
Yoder (1973), drawn at the first appearance
of akermanite, between 1 and 5.75 kbar is
about 58°C/kbar. The reaction takes place
through a range of temperatures 50°-70°C
wide at a given pressure, presumably be-
cause of possible solid solution of merwinite
in akermanite, diopside in merwinite
(Schairer et aiy 1967; Yoder, 1973), and
minor substitution of Ca by Mg in calcite.
The akermanite + diopside + calcite region
was not observed in the present study when
akermanite + calcite and diopside + 2 cal-
cite compositions were used. The newly
crystallized akermanite, however, contains
tiny inclusions of (relict) diopside that are
separated from calcite by the akermanite
host. The diopside + calcite <=> akermanite
+ CO2 reaction curve, now bracketed by
means of akermanite + calcite and diopside
+ 2 calcite compositions, has a revised
slope of about 62°C/kbar. It terminates at
the invariant point I\, which was found to
be located at5.4±0.2kbar and 1215°±5°C
with akermanite + diopside + calcite +

liquid + CO2 in equilibrium.

The invariant point I\ is the locus of four
other reactions occurring clockwise:

[CO2] akermanite + calcite + diopside
<=> liquid,

[ak] calcite + diopside

<=> liquid + CO2,

[cc] diopside + liquid

<=> akermanite + CO2, and

[di] akermanite -1- calcite + CO2 <=> liquid.

Separation of reaction [C02] from reaction
[ak] was not possible because the liquid
field may be located very close to or even
on the diopside + calcite join. Above the
temperature of reaction [ak], diopside +
liquid + C02 will reach the diopside + 2
calcite composition (= compositional sin-
gularity), and any further increase in tem-
perature will move the liquid -1- C02 tie
lines along diopside + calcite toward the
join akermanite -1- CO2, that is, becoming
compositionally equivalent to diopside + 1
calcite. That particular composition is
reached in S2 at 13 10°C and about 5.3 kbar
located on reaction [cc], diopside + calcite
= akermanite + CO2, from which the two
restricted reactions diopside + calcite =
liquid (higher pressure limb) and liquid =
akermanite -1- CO2 (lower pressure limb)
will occur.

Merwinite Relations With CO2 and H2O

Experiments on the stability of
merwinite in the presence of C02 + H2O
were conducted at 1 and 3 kbar (Figs. 50A
and 50B). The merwinite + V stability field

GEOPHYSICAL LABORATORY

1300

1200

1100

°- 1000
0

| 900
CD

i® 800

700

600

P=1 kbar

(1)KUSHIRO&YODER,1964

(2) WALTER, 1965

(3)YODER,1973

(4) HUCKENHOLZ etal., 1990

0.0 0.1 0.2 0.3 0.4 0.5 0.6 0.7 0.8 0.9 1.0

XcOo

1300

_ . - 995'

1195*
1170'

1090'
1065'

Fig. 50a and 50b. Isobaric temperature -X(<X>2) [X((X>2) = CO2ACO2 + H2O)] diagram for 1 kbar (Fig.
50A) and 3 kbar (Fig. 50B). Abbreviations for phases are: mo, monticellite (CaMgSi04); per, periclase
(MgO); fo, forsterite (Mg2Si04) spur, spurrite [Ca5Si208(C03)]; wo, wollastonite (CaSi03); geh,
gehlenite (Ca2Al2Si07); gro, grossular (Ca3Al2Si30i2); cor, corundum (AI2O3); mei, meionite
(Ca3 Al6Si6024*CaC03); V, fluid; other abbreviations and symbols as in Fig. 49. Letters on half-shaded
symbols or symbols with dots refer to experiments on mixtures F, D, E, Jj, and J6- The solidus of these
compositions (except for E) refers to ak + di + CO2 (above 1 1 50°C). Light curves and light dashed curves
are decarbonation reactions (extrapolated) from Walter (1963a.b; 1965) h to/6 are not fully illustrated.
The geh + V reaction is from Huckenholz et al. (1990).

CARNEGIE INSTITUTION

decreases with increasing pressure when
the C02 fluid is diluted with H2O. At
AXCO2) = 1.00, akermanite + calcite =
merwinite + V are restricted to a tempera-
ture of 1 1 20°C at 1 kbar and 1 1 82°C at 1 .3
kbar (isobaric invariant point I2). From I2
the reaction akermanite + calcite <=>
merwinite + V shifts to lower AXCO2) when
the pressure increases (light dashed curves
in Fig. 50A and 50B). At 810°C and at 1
kbar the reaction spurrite -1- monticellite <=>
merwinite + calcite (Walter, 1965), well
displayed at Crestmore, California
(Burnham, 1959), must intersect with
akermanite + calcite <=> merwinite + V at
about X(C02) = 0.10 in Fig. 50A The
resulting invariant point labelled I5 be-
comes the locus of the reaction of spurrite
+ akermanite + monticellite <=> merwinite
+ V and akermanite + calcite <=> spurrite +
monticellite + V (not shown in Fig. 50A).
With decreasing temperatures spurrite +
akermanite + monticellite <=> merwinite +
V intersects monticellite + wollastonite <=>
akermanite at a pressure of 1 kbar and a
temperature as low as 700°C and atX(C02)
= 0.07 (isobaric invariant point Ie).

Merwinite + calcite + V will melt at 1
kbar over almost the entire range of X(C02)
displaying merwinite + calcite + liquid as
well as the merwinite -1- liquid + V assem-
blage. Increasing pressure shifts the
merwinite + calcite + V solidus toward
lower X(C02) and lower temperatures as
well. At 3 kbar, merwinite + calcite + V will
melt atX(C02) < 0.15 and temperatures <
1150°C. At X(C02) > 0.15, akermanite +
calcite + V forms the subsolidus assem-
blage, which melts between 1150° and
1 200°C to akermanite + calcite + liquid and
akermanite + liquid + V assemblages. The
X(C02) of the liquid coexistent with fluid

was not determined during the course of the
present study.

Melting experiments on silicate + H2O
+ CO2 systems (Mysen, 1975; Rosenhauer
and Eggler, 1975; Kadik and Eggler, 1975)
demonstrate, however, the preference of
H2O solubility over CO2 in silicate melts
and, therefore, X(C02) of the liquid of
about 0.05 was assumed for the tempera-
ture and pressure conditions studied. The
diopside + calcite = akermanite + CO2
reaction was determined for 3 kbar at
X(C02) of 1 .00, 0.50 and 0.05 (Huckenholz
et ai, 1990; this study) and found to be
1065°, 1010°, and 740°C, respectively. Cal-
culations for X(C02) (Helgeson et ai,
1978; Holloway; 1977) at 0.95, 0.10, and
0.05 yielded temperatures of 1059°, 846°,
and 803 °C. Experimental data for 1 kbar
are available only for X(C02) at 1 .00 and
950°C and for X(C02) at 0.5 and 895°C
(Huckenholz etal., 1990; this study). They
were calculated for X(C02) = 0.10 and
0.20, and found to be about 735° and 770°C,
respectively.

The diopside + calcite <=> akermanite +
V reaction is intersected by the fluid- and
calcite-absent reaction of diopside +
monticellite = forsterite + akermanite on
the basis of the experimental data of Kushiro
and Yoder (1964) and of Yoder (1973). The
intersection will occur at 1 kbar at 880°C
and about X(C02) = 0.45 and at 3 kbar at
about 920°C and X(C02) = 0.20 ± 0.02,
resulting in the isobaric invariant point
labeled I4. The validity of the diopside +
monticellite = forsterite + akermanite reac-
tion is questioned by Helgeson etal. ( 1 978),
who calculated reaction temperatures of
about 1 300°C for 1 kbar and about 1 350°C
for 3 kbar. It should be noted, however, that
the anhydrous melting of diopside +

GEOPHYSICAL LABORATORY

akermanite + forsterite assemblages will
occur at 1357°C at 1 arm (Ricker and
Osborn, 1954). A hydrous fluid phase
involved in the melting will lower the
solidus considerably, well below the tem-
perature of the solid- solid reaction as cal-
culated by Helgeson et al. (1978). The
diopside + calcite <=> akermanite + CO2
reaction will terminate at the CO2- and
calcite-absent reaction of monticellite +
wollastonite <=> akermanite (Yoder, 1973;
Huckenholz, 1990; this study). The result-
ing isobaric invariant point I4 for 1 kbar
occurs at 700°C and atX(C02) = 0.1; for 3
kbar, it is found to be at 720°C and at
X(C02) = 0.05 ± 0.02.

Discussion

With the experimental data at hand, it
has been demonstrated that merwinite +
CO2 is stable over a wide range of tempera-
ture but is restricted to pressures below
about 1.5 kbar. Merwinite + V, however,
may also form atX(C02) as low as 0.05 and
at temperatures as low as 700°C from
spurrite + monticellite + akermanite and
from spurrite -1- monticellite + wollastonite
assemblages in H20-rich aureoles of car-
bonaceous rocks around basic intrusives.
At pressures below 1 .3 kbar merwinite +
calcite +V and merwinite + V assemblages
reach the solidus between 1 1 50° and 1 200°C
and will melt at any X(C02) up to 1 .0. At
pressures above 1 .3 kbar the merwinite + V
stability shifts toward lower X(C02) and
akermanite + calcite + V assemblages reach
the solidus and are subject to melting. At 3
kbar, these same relationships occur at
1 1 50°- 1 200°C and X(C02) > 0. 1 5 .

High-temperature phase assemblages
occur as rock inclusions in the critical zone

of the Bushveld Complex:

( 1 ) merwinite + monticellite + melilite
(equivalent to mixture E),

(2) calcite + periclase + monticellite,
and

(3) forsterite + periclase +
monticellite.

A different set of high-temperature phase
assemblages occur as rock inclusions in the
marginal zone:

(4) akermanite + monticellite + cal-
cite (equivalent to mixture D),

(5) calcite + forsterite + monticellite,

(6) akermanite + diopside +
monticellite (equivalent to inclusions
Jl and J6), and

(7) diopside -1- forsterite + monticellite.
According to Wallmach et al. (1989), tem-
peratures as high as 1300° and 1200°C,
respectively, for estimated pressures at
about 1 kbar (0.6 to 1 . 1 kbar) and 2 kbar
(1.1 to 2.4 kbar) are deduced. Experimen-
tal results obtained on merwinite + V as-
semblages (mixtures A, B, C, and Y),
merwinite + monticellite + melilite assem-
blages (mixture E), akermanite +
monticellite + calcite assemblages (mix-
ture D) as well as on the rocks (Ji and J6)
containing akermanite + monticellite + di-
opside are clearly above their solidi when
treated at 1200°C and at pressures of 1 and
3 kbar, respectively, with X(C02) between
0.3 to 1.0. Thus, the deduced temperatures
of Wallmach et al. (1989) appear to be
excessive.

References

Agrell, S. O., Poly thermal metamorphism of lime-
stones at Kilchoan, Ardnamurchan, Mineral.
Mag., 34, (Tilley vol.) 1-15, 1965.

Bowen, N. L., Progressive metamorphism of sili-
ceous limestone and dolomite, J. Geoi, 48,
225-274, 1940.

CARNEGIE INSTITUTION

Burnham, C. W., Contact metamorphism of mag-
nesian limestone at Crestmore, California,
GeoL Soc, Amer. Bull, 70, 879-920, 1959.
Helgeson, H. D., J. M. Delany, H. W. Nesbitt, and
D. K. Bird, Summary and critique of the ther-
modynamic properties of rock-forming min-
erals, Amer. J. Sci., 278-A, 1-229, 1978.
Holloway, J. R., Fugacity and activity of molecu-
lar species in supercritical fluids, in Thermo-
dynamics in Geology, D. Fraser, ed., 161-181,
1977.
Huckenholz, H. G., A. Wassermann, and K. T.
Fehr, Stability and phase relations of gehlenite-
akermanite solid solutions in the presence of a
H2O-CO2 fluid, Third International Sympo-
sium Experimental Mineralogy , Petrology, and
Geochemistry, p. 81, Edinburgh, 1990.
Joesten, R. C, Pseudomorphic replacement of
melilite by idocrase in a zoned calc-silicate
skarn, Christmas Mountains, Big Bend Re-
gion, Texas, Amer. Mineral., 59, 694-699,
1974.
Joubert, J., Gemetamorfoseerde karbonaatins
luitsels in die bosone van die
Bosveldstollingskompleks, was van
Roossenekel, Transvaal, M.Sc. thesis, Univ.
Pretoria, Hillcrest, South Africa, 1976.
Kadik, A. A., and D. H. Eggler, Melt- vapor rela-
tions on the join NaAlSi30s-H20-C02,
Carnegie Instn. Washington Year Book, 74,
479-484, 1975.
Kushiro, I., and H. S. Yoder, Jr., Stability field of
akermanite, Carnegie Instn. Washington Year
Book, 63, 84-86, 1964.
Larsen, E. S., and W. F. Foshag, Merwinite, a new
mineral from the contact zone at Crestmore,
California, Amer. Mineral., 6, 143-148, 1921.
Mysen, B. O., Stability of volatiles in silicate
melts at high pressure and temperature,
Carnegie Instn. Washington Year Book, 74,
454-478, 1975.
Ricker, R. W., and E. F. Osborn, Additional phase
equilibrium data for the system CaO-Mg02-
Si02, /. Amer. Ceram. Soc, 37, 133-139,
1954.
Rosenhauer, M., and D. H. Eggler, Solubility of
H2O and CO2 in diopside melt, Carnegie
Instn. Washington Year Book, 74, 474-479,
1975.
Schairer, J. F., H. S. Yoder, Jr., and C. E. Tilley,
Behavior of synthetic melilites in the join
gehlenite - soda melilite - akermanite, Carnegie
Instn. Washington Year Book, 65, 217-226,
1967.
Shmulovich, K. I., Stability of merwinite in the
system CaO-MgO-Si02-C02,D^/./4c^.5d,
USSR, Earth Sci. Sect., 184, 125-127, 1969.
Smith, J. V., X-ray emission microanalyses of
rock-forming minerals, VI. Clinopyroxenes
near the diopside-hedenbergite join, /. GeoL,

74, 463-477, 1966.

Tilley, C E., On larnite (calcium orthosilicate, a
new mineral) and its associated minerals from
the limestone xenoliths in the eastern Bushveld
Complex, Canad. Mineral.,27, 509-523, 1989.

Walter, L. S., Experimental studies on Bowen's
decarbonation series, I: P-T univariant equi-
libria of the "monticellite" and "akermanite"
reactions, Amer. J. Sci., 261, 488-500, 1963a.

Walter, L. S., Experimental studies on Bowen's
decarbonation series II: P-T univariant equi-
libria of the reaction: forsterite + calcite =
monticellite + periclase + CO2, Amer. J. Sci.,
261, 173-179, 1963b.

Walter, L. S., Experimental studies on Bowen's
decarbonation series III: P-T univariant equi-
libria of the reaction spurrite + monticellite =
merwinite + calcite and analyses found at
Crestmore, California, Amer. J. Sci., 263, 64-
77, 1965.

Yoder, H. S., Jr., Melilite stability and paragen-
esis, Fortschr. Mineral, 50, 140-173, 1973.

Yoder, H. S., Jr., Relationship of melilite-
bearing rocks to kimberlite. A preliminary
report on the system akermanite-C02, Phys.
Chem. Earth, 9, 883-894, 1975.

The System MG2S1O4-FE2S1O4
at Low Pressure

Hiroko Nagahara, Ikuo Kushiro* and
Bjorn O. Mysen

Gas -solid relationships are important
when we consider condensation, evapora-
tion, and fractionation of the solar nebula,
especially in regard to bulk composition of
the Earth and the terrestrial planets. Gas-
solid relationships of minerals are different
from those between liquid-gas, and ther-
modynamic data are insufficient to con-
struct gas-solid phase diagrams. Mysen
and Kushiro (1988) and Kushiro and Mysen
(1991) measured vapor pressures of MgO,

* Geological Institute, University of Tokyo, Hongo,
Tokyo 113, Japan

GEOPHYSICAL LABORATORY

Si02, forsterite, and enstatite, and studied
the phase relations of the MgO-Si02 sys-
tem. In this study, we have measured the
vapor pressure of fayalite, and based on
those results along with those for forsterite,
the gas-solid phase relationships of the
olivine system as a function of the Fe/
(Mg+Fe) ratio are proposed.

Vapor pressure measurements of fayalite
were made using a Knudsen cell method
similar to that described by Mysen and
Kushiro (1988). The starting material is a
single crystal of fayalite, about 2x2x5 mm
in size, synthesized with the Czochralski-
pulling method by H. Mori of the Univer-
sity of Tokyo. The crystal was powdered
(1-10 Jim), and several mg of the sample
was placed in a molybdenum capsule with
two 2-mm holes drilled in sides. The ex-
periments were carried out in vacuum fur-
naces in the University of Tokyo and the
Geophysical Laboratory; the two furnace
designs are nearly identical in size and
design (Mysen and Kushiro, 1 988). Samples
were heated at a rate of 15-20°/min from
room temperature to experimental tem-
peratures, which ranged from 1050°C to
1 175°C. Run durations ranged from 4 days
at 1175°C to 12 days at 1075°C. Total
pressure of the vacuum chamber was 4.0 x
10-7 torr (5.3 x 10"1(> bar) to 6.0 x 10-7 torr
(7.9x10-9 bar).

Experiments in molybdenum capsules
were at the oxygen fugacity (fo2) of the
M0-M0O2 buffer to ensure that fayalite is
stable. The M0-M0O2 buffer is about 1.5
orders of magnitude higher than the iron-
wiistite (IW) buffer, and 1 to 1.5 orders of
magnitude below the quartz-fayalite-mag-
netite (QFM) buffer (Mysen and Kushiro,

1988). Measured weight loss (1-14 %) was
calibrated against vapor pressure by the
equation

P - 1

1 m —

Ac dt

2kRT
M

where Pm is the vapor pressure of a sub-
stance, A is the area of the orifice of the
capsule, c is the clausing factor, dw is the
weight loss, dt is duration, M is the
molecular weight of the effusing vapor, 7is
the absolute temperature, and R is the gas
constant (Paule and Margrave, 1967). The
clausing factor for the cell has previously
been determined for Cu and Ag by Mysen
and Kushiro (1988).

The residue of partial evaporation re-
mained fayalite, indicating that fayalite
evaporates congruently. Since forsterite
evaporates congruently and intermediate
olivine evaporates stoichiometrically
(Mysen and Kushiro, 1988; Nagahara et
al.y 1988), it is clear that olivine evaporates
stoichiometrically regardless of the Fe/Mg
ratio. Fayalite heated at 1175 °C melted,
but fayalite heated at 1 160°C did not, sug-
gesting that the melting point is about
1 170°C and 2 x 10-8 bar. This is about 35°C
lower than that at 1 bar (Bo wen and Schairer,
1935). Lower melting temperatures rela-
tive to melting temperature at 1 bar have
also been found for Si02 and Mg2Si04 at
low pressures (by 100° and 200°C, respec-
tively) (Mysen and Kushiro, 1988). Mysen
and Kushiro (1988) further demonstrated
the presence of a three-phase region be-
tween that of solid and gas and that of liquid
and gas. This three phase region (forsterite
+ liquid + vapor) implies that the system

CARNEGIE INSTITUTION

1040

1080

1120
T (<>C)

1160

1200

Fig. 5 1 . Temperature and vapor pressure relation-
ship of fayalite. Error bar represents weight mea-
surement uncertainty.

can not be described as a binary (MgO-
S1O2). Therefore, phase relations where
liquid is present will not be discussed in the
present study.

The experimentally determined tem-
perature and vapor pressure (Pv) relation-
ships for Fe2Si04 are summarized in Fig.
51. The relationship is further shown in

-12 ■

-16

-20

-24

1/TxlO4 (1/K)

Fig. 52. Arrhenius plot for vaporpressure of fayalite
(this work) and forsterite, MgO, and Si02 (Mysen
and Kushiro, 1988). Linear regression line from
the fayalite data gives In />v=(-608±60)/Rr +
(273±9)/R and r = 0.9976.

Fig. 52 together with the l/T vs. In Pv of
Mg2Si04, MgO, and Si02- The linear re-
gression curve from the data points can be
expressed in the equation for evaporation

In Pv = :Mv + ASv
RT R

(2)

where A//v and ASV are enthalpy and en-
tropy of evaporation, respectively. From
the linear regression, A//v is 608 ± 60 (1 o)
(kJ/mol) and ASV is 273 ± 9 (J/K-mol).
These values are similar to those for
forsterite (640 ± 36 and 210±54, respec-
tively) at the same/02 (Mysen and Kushiro,
1988).

The gas-solid phase diagram was drawn
by using the enthalpies and entropies for
evaporation of forsterite and fayalite (Fig.
53). In order to draw the phase diagram,
two assumptions were made: (1) chemical
equilibrium is achieved between crystals
and gas in a Knudsen cell, and (2) both
olivine and gas are ideal solutions. The
assumption (1) can be valid. The residues
are sintered homogeneous fayalite crystals
regardless of experimental duration. Ac-
cordingly, chemical compositions of coex-
isting gas and solid have been uniform. The
assumption (2), ideality of the olivine solid
solution system, has been shown by many
investigators (i. e., Wood and Kleppa, 1981).
Gas can be treated as ideal.

With the assumptions made above, mol
fractions of the fayalite component in gas
and solid at a given pressure are shown by
the following equations

GEOPHYSICAL LABORATORY

1800

-" 1 • r

^ 1600

■*->
(0

0)
Q.

£ 1400

1200

1800

1600

■*-•

1400

1200

GAS

OLIVINE s.s.+ GAS

I0"8bar

J I l_

■ I I 1_

20 40 60 80 100

"> 1 " r

T r

10"10 bar

GAS

_. 1 i_

1800

^ 1600

1200

-i 1 «"

GAS

10 bar
j i i k

20 40 60 80 100

Fe/(Mg+Fe)x100

20 40 60 80 100

Fe/(Mg+Fe)x100

Fig. 53. Gas-solid phase diagrams of the olivine
system at 10"8, 10A and 10r° bar.

ln*- =
x

and

'_ AHFi

\T0 ~fl

\-x_ AHFo
1 -x

(3)

(4)

where x and xy are the Fe/(Fe+Mg) ratios of
gas and solid, respectively, and T0and Tq
are vaporus temperatures for fayalite and
forsterite, respectively.

The calculated vaporus and solidus
curves are shown in Fig. 53. The conspicu-
ous feature of the figure is that the binary
loop is quite flat. With decreasing pressure,
the vaporus temperatures for both forsterite
and fayalite become lower and the loop
becomes more flat. The figure sows that the
compositional difference between coexist-
ing solid and gas is extremely large com-
pared to that between solid and liquid at L
bar (Bo wen and Schairer, 1935). There is a
very narrow temperature interval over
which Fe -bearing olivine coexists with gas.
Thus, olivine should have become forsterite

CARNEGIE INSTITUTION

regardless of the primary composition when
it was heated at subvaporus temperatures.
Formation of forsterite would have been
kinetically suppressed when equilibrium
between gas and solid was not achieved.
Hashimoto (1990) showed that forsterite
evaporates very slowly at a free evapora-
tion condition (at a rate one-tenth of that in
equilibrium); that is, if the generated gas
was removed from the system and did not
equilibrate with the solid, formation of
forsterite would be suppressed. Another
possible factor preventing formation of
forsterite is cation diffusion in solid
forsterite, which depends on temperature,
heating duration, and the grain size of oli-
vine. Although the Mg-Fe inter-diffusion
coefficient in olivine is largest among any
other elemental diffusion coefficients in
olivine and those known in any other sili-
cate minerals (Freer, 1981), short heating
and/or large grain size can be rate-limiting
factors for formation of forsterite as partial
evaporation residue. If these processes were
not effective, olivine should have become
forsterite quickly by heating at subvaporus
temperatures.

The phase diagram is not directly appli-
cable to evaporation in the solar nebula;
effects of other components, pressure, and
foi should be evaluated. Other components,
such as Al and Ca, would affect the abso-
lute temperature of vaporus and solidus,
but would not change significantly the shape
of the diagram because of much smaller
abundance in the solar nebula of these
components compared with Si, Mg, and Fe
(Anders and Ebihara, 1982). The pressure
range in the present experiments is just

applicable to the solar nebula. Elemental
abundances of Mg and Si in the solar sys-
tem are about 4 orders of magnitude smaller
than that of H, and the total pressure at the
midplane of about 3 A.U. has been gener-
ally calculated to be between 10-3 and
10-5 bar (Cameron, 1985; Morfill et al.,
1985). The approximate partial pressure
for olivine component is thus 107 to 10-9
bar, which well agree with the pressure
range in the present work. Oxygen fugacity
condition of the present work is much more
oxidizing than that estimated for the solar
nebula, based on the elemental abundances
of the solar system. However, oxygen fu-
gacity as high as the present work has been
proposed recently for the formation of vari-
ous chondritic components in the solar
nebula (Fegley and Palme, 1985; Palme
andFegley, 1990; Weinbruclmtf/., 1990).
The present result can be, thus, applied
almost directly to the conditions for forma-
tion of chondritic components in the solar
nebula.

References

Anders, E., and M. Ebihara, Solar-system abun-
dances of the elements. Geochim. Cosmochim.
Acta, 46, 2363-2380, 1982.

Bowen, N. L. and Schairer, J. F., The system
MgO-FeO-Si02, Amer. Jour. Sci., ser. 5, 29,
151-217, 1935.

Cameron, A. G. W., Formation and evolution of
the primitive solar nebula, in Protostar and
Planets II, D. C. Black and M. S. Matthews, eds.,
Univ. Arizona Press, Tucson, Arizona, pp. 1073-
1099, 1985.

Fegley, B., Jr. and H. Palme, Evidence for oxidiz-
ing conditions in the solar nebula from Mo and
W depletion in refractory inclusions in carbon-
aceous chondrites, Earth Planet. Sci. Lett., 72,
311-326,1985.

GEOPHYSICAL LABORATORY

Freer, R., Diffusion in silicate minerals and glasses:
A data digest and guide to the literature. Contrib.
Mineral Petrol, 76, 440-454, 1981.
Hashimoto, A. Evaporation kinetics of forsterite
and implications for the early solar nebula.
Nature, 347, 53-55, 1990.
Kushiro, I. and B. O. Mysen, in Progress in
Metamorphic and Magmatic Petrology , L. L.
Perchuked, Cambridge Univ. Press, Cambridge,
England, 411-433, 1991.
Morfill, G. E., W. Tscharnuter, and H. Volk, in
Protostar and Planets II, D. C. Black and M. S.
Matthews, eds., Univ. Arizona Press, Tucson,
Arizona, pp. 493-533, 1985.
Mysen, B. O. and I. Kushiro, Condensation,
evaporation, melting, and crystallization in the
primitive solar nebula: Experimental data in the
system MgO-Si02-H2 to 1.0xl0-9barand 1870°C
with variable oxygen fugacity, Amer. Miner.,
73, 1-19, 1988.
Nagahara, H., I. Kushiro, B. O. Mysen, and H.
Mori, Experimental vaporization and condensa-
tion of olivine solid solution, Nature, 331, 516-
518, 1988.
Palme, H. and B. Fegley, Jr., High temperature
condensation of iron-rich olivine in the solar
nebula, Earth Planet. Sci. Lett., 101, 180-195,
1990.
Paule, R. C. and J, L, Margrave, Free-evaporation
and effusion techniques, in The Characteriza-
tion of High Temperature Vapors, J. L. Margraves
ed., Wiley, New York, 130-151, 1967.
Weinbruch, S., H. Palme, W. F. Muller, and A. El
Goresy, FeO-rich rims and veins in Allende
forsterite: Evidence for high temperature con-
densation at oxidizing conditions, Meteoritics,
25, 115-125, 1990.
Wood, B. J. and O. J. Kleppa, Thermochemistry of
forsterite-fayalite olivine solutions, Geochim.
Cosmochim. Acta, 45, 529-534, 1981.

Fe3+, Mg Order-Disorder in Heated

MGFE2O4: A Powder XRD

and 57Fe Mossbauer Study

H. St. C. O'Neill; H. Annersten,**
and D. Virgo

The two extreme cation distributions in
the spinel structure are the so-called nor-
mal configuration A [Bii O4 and the inverse
configuration B[AB]04 where the [ ] refer
to the octahedrally coordinated cations and
the remaining cations are in tetrahedral
coordination. Disordered configurations of
both these extreme arrangements can be
represented as A\.x BX[AX #2-x]04. The
degree of disorder can also be discussed in
terms of the inversion parameter x, defined
as the fraction of B cations occupying the
tetrahedral sites.

A general thermodynamic model of the
cation distribution in spinels has been pro-
posed by O'Neill and Navrotsky (1983,
1984). The basic tenet of this model is that
the equilibrium cation distribution is re-
lated to the free energy of disordering in the
following way,

•RT\A—^ U^SL

\(l-x)(2-x)j \ dx Itj,,

where AGd is the change in the non-con-
figurational free energy of disordering.

In their model, the free energy term
AGd was shown to consist of an enthalpy

*Bayerisches Geoinstitut, Universitat Bayreuth,
Germany

** Department of Mineralogy and Petrology, Insti-
tute of Geology, University of Uppsala, Sweden

CARNEGIE INSTITUTION

term, AHd which takes a quadratic depen-
dence on x (AHd = OUC+ Pjc^). There was
also a non configurational entropy term
ASd- The non-linear nature of the enthalpy
term implies that cation site preference will
depend on the degree of inversion of the
spinel into which it is substituting. Previ-
ously it had been proposed (Navrotsky and
Kleppa, 1967) that the site preference en-
thalpies in spinels were independent of
both temperature and the degree of order in
the spinel structure.

The experimental basis for the proposed
model of non-linear enthalpy of disorder-
ing in spinel solid solutions is limited
(O'Neill and Navrotsky, 1983, Nell et aL,
1989). Therefore, the present ordering and
disordering experiments on MgFe204 have
been specifically designed to test the above
thermodynamic model. Magnesio-ferrite
is an ideal composition in which to carry
out such tests, since it shows a relatively
large change in equilibrium cation distri-
bution over a wide temperature range, it is
stable in air, and the degree of inversion can
be determined by several techniques such
as x-ray refinement, and 57Fe Mossbauer
spectroscopy. Data in the literature are
viewed as unsatisfactory, in view of the
possibility either that the samples studied
were not stoichiometric or that the cation
site populations do not correspond to the
annealing temperatures.

Experimental

MgFe204 was synthesized in air using
a sodium tungstate flux. The oxide compo-

nents consisting of 2 g MgO, 4 g Fe203, 20
g Na2W04, and 2 g WO3 were melted in a
Pt crucible at 1260 C and then cooled at
6°C per hour to either 950°C (first batch) or
900°C (second batch), at which tempera-
tures the melt was crystallized for approxi-
mately 12 hours before the final cooling to
room temperature. The sodium tungstate
flux was dissolved in warm water. Excess
MgO was removed in dilute nitric acid.
The MgFe204 run products consisted of
euhedral, octahedra of reddish-brown
spinel, -10 pxn in size.

The acid cleaned MgFe204 was
analysed by ICP and by two different elec-
tron microprobe systems, one at the
Bayerishes Geoinstitut and the other at the
University of Uppsala. The MgO/Fe203

values were respectively 1.014 ±
0.019,1.014±0.009andl.028±0.015.All

three sets of analyses indicate that MgFe204
is stoichiometric within the analytical un-
certainty (ie. within 2 standard deviations).
Ideal stoichiometry is assumed in the fol-
lowing discussion of the x-ray and 57Fe
Mossbauer results.

Both disordering and ordering experi-
ments were carried out in vertical quench
furnaces in the range 400-1250°C. by heat-
ing 40-100 mg of the starting material in
platinum capsules welded at one end and
crimped at the other end. Temperatures
were controlled to ±0.5 °C, and run dura-
tions were up to 50 days. The heated
samples were quenched in water. Effective
quench times calculated by extrapolation
of the rate studies were of the order of a few
milliseconds.

GEOPHYSICAL LABORATORY

Lattice Parameter Measurements and
Powder XRD Structural Refinements

Lattice parameters were measured us-
ing CoA'ai radiation on a STOE focusing
diffractomer with a curved Ge monochro-
mator. A position-sensitive detector of 8°
20 width and 0.5° data interval was used to
collect the diffraction data over the range
40-120° 2£ NBSSi (0 = 5.43087) was used
as an internal standard. Peak positions
were determined by fitting the peak profile,
and the observed 20 's were adjusted by
using a linear correction based on the six
main Si peaks in the two regions of interest.
The spinel lattice parameter a was then
determined by a weighted least squares
refinement of the positions of the ten most
intense spinel peaks. The mean internal
estimate of one standard deviation of a is
calculated to be ± 0.00015. An external
estimate, based on determinations of a for
samples equilibrated at 600, 650 and 700° C
and heated for times that exceed that re-
quired for equilibrium gives estimates in
theuncertainty of0of±O.OOOll toO.00013.

The XRD structural refinements were
measured on data using the same
diffractometer but with Mo/foci radiation
and over the range in 20 from 7 to 80°. Data
were collected from five consecutive scans
and the scans were summed. Crystal struc-
ture refinements were made using two con-
trasting methods which are more fully dis-
cussed by O'Neill et al. (1991a). Here,
results for full profile Rietveld-type refine-
ments using the program DBW (Wiles and
Young, 1987) are considered (see also
O'Neill et al, 1991b) The structure was
refined in space group Fd3m with 8(a)

andl6(d) cation sites and 32(e) oxygen
sites all fully occupied. The refined quanti-
ties were (1) a scale factor, (2) the Pearson
VII profile shape parameters in which the
component m could vary as m = a + b/20 ,
(3) an asymmetry parameter for peaks with
2q < 34°, (4) a 20 zero parameter, (5) Peak
halfwidth function of the form FWHM2 =
UtmO + VtmO + W, (6) the unit cell
constant a (7) the oxygen positional param-
eter m , (8) the inversion parameter, x, and
(9) isotropic temperature factors for either
cation sites or oxygen, or both. Refine-
ments were also made for the cases where
the raw data background was either in-
cluded or subtracted. Bragg and total pat-
tern residuals for the equilibrated samples
were in the ranges 1 .97-3.63 and 2.89-4.57,
respectively. There is no systematic corre-
lation between these residuals and the de-
gree of inversion.

57Fe Mossbauer Spectroscopy

Conventional transmission spectra were
collected in a 512 multi-channel analyzer
operated in conjunction with a constant
accelerator electro-mechanical drive unit.
57Co in Rh was used as a source. A calibra-
tion spectrum for natural iron at room tem-
perature was simultaneously recorded at
the other end of the vibrator unit. The two
mirror-symmetric spectra typically con-
tained 0.6-0.8 x 10^ counts/channel, and
the counts/channel for each half of the
spectral data were averaged before analy-
sis by a least-squares fitting program.

The room temperature spectra of
magnesio-ferrite (Neel point around 600

CARNEGIE INSTITUTION

100.0

o 99.0

8 98.0

§ 97.0

w ~~ ~

<i> 96.0

95.0 -

•5 0 5

Velocity, mm/s

o
w
-Q

cz
CO
c
o

CD
DC

100.0

o 99.0

98.0 -

E 97.0 -

96.0
95.0

-10

0
Velocity, mm/s

Fig.54. 57Fe Mossbauer spectra of stoichiometric MgFe204 measured in an
external field of 4.5 Tesla. The absorption patterns labeled A and B refer to Fe3+
in the/4 and£-sites, respectively. Upper spectrum is MgFe204 heated at 1050°C,
35 mins., Absorber temperature = 171 K. Lower spectrum is MgFe2C>4 heated
500° C, 8 days, Absorber temperature = 189 K.

K) indicates that there is almost complete
overlap of the magnetically split patterns
due to Fe3+ in the tetrahedrally and octahe-
drally coordinated sites. Accordingly, the

spectral data were collected in the presence
of an externally applied magnetic field
using a superconducting magnet, cooled to
4.2 K to produce a magnetic field of 4.5

GEOPHYSICAL LABORATORY

Tesla. In these experiments, the external
magnetic field is orientated parallel to the
propagation of the gamma ray and thus the
intensity of the transitions AMi =0 vanish;
in addition, the magnetic hyperfine field at
the tetrahedral site increases while that at
the octahedral site decreases. Thus, each
six-line pattern of the 298 K spectra of
MgFe204 is reduced to four lines in the
presence of the external field (cf Fig. 54).
The absorbers were made by mixing
powdered samples of equilibrated and
quenched MgFe204 with -100 mg of
thermo-setting plastic powder, (cf Virgo
and Hafner, 1969). The absorber density
was 7 mg Fe/cm2- The absorber was cooled
in the cryostat to temperatures in the range
1 2- 1 7 1 K; this procedure minimized differ-
ences in the recoil-free fraction at the A and
B sites.

Least squares fitting of the spectral data,
carried out using lines of Lorentzian shape,
incorporated equal half- widths of each four-
line pattern, and the intensity ratios were
independent of whether the line intensities
were unconstrained or whether the mag-
netically split lines 1, 3, 4, and 6 were
constrained in the ratio 3:1:1:3. It is evi-
dent from Fig. 54 that the absorption due to
Fe3+ in the octahedrally coordinated sites
is significant broader compared to that for
Fe3+ in the tetrahedrally coordinated sites.
It is also evident from Fig. 54 that the half-
width of the Z?-site absorption increases
with increasing disorder. For the sample,
heat treated at 1050°C, four hyperfine split
patterns were refined in the fit of Fe3+(#-
site) whereas only three such patterns were
required for the remaining heat-treated and
equilibrated samples in order to obtain sta-

8.41

<
._- 8.40

CD
CD

cfl

co 8.39

CD
O

8.38-

8.37

I I I I I I I I
_ MgFe204

■ ■ ■ ■

■
■

— ■ —
■

■

_ ■ _

i i i i i i i i

400

600 800 1000 1200

Temperature of anneal, °C

Fig. 55. Lattice parameter (a) measurements on
heat-treated and quenched samples of stoichio-
metric MgFe2C>4.

tistically meaningful fits. For all values of
x only a single hyperfine pattern was fitted
to the absorption due to Fe3+ in the tetrahe-
dral site.

Approach to Equilibrium

The lattice parameters were used to
monitor the approach to the equilibrium
cation distribution in both disordering and
ordering experiments. It can be shown
from the results of this study that a varia-
tion of ±0.0001 in a corresponds to a
variation in x of about ±0.001. The cell
constants for samples annealed and then
quenched at 50° intervals in the range 450-
1250°C are shown plotted in Fig. 55. Equi-
librium is achieved on a time scale of about
30 days at 450°C and less than five minutes
at temperatures of 700°C and above. For
equilibrium compositions in the range 550-
700° C, equilibrium was verified by rever-
sal experiments whereby the equilibrium
values of x, at fixed temperature, were
determined with starting material that was

CARNEGIE INSTITUTION

both disordered and ordered with respect to
the equilibrium value (O'Neil , 1991c).

The data plotted in Fig . 5 5 show a smooth
increase in a as a function of temperature,
although there is only a small increase in a
above 1050°C. Itisofconcernatthesehigh
temperature that the rate of ordering is so
fast in the high temperature experiments
that the equilibrium distribution of Mg and
Fe3+ is not quenched-in. There is also the
possibility that there are deviations from
the stoichoimetric composition at high tem-
perature. While both those possibilities are
discussed in more detail by O'Neill et al.
(1991b), it is significant that the thermody-
namic model to be discussed below and
established from a fit to the cation distribu-
tion on samples annealed at temperatures
of less than 1000°C does, in fact, predict
that there will only be comparatively small
changes in ao at temperatures above 1 100°C.

of the B -values with the degree of inver-
sion.

The values of x are shown plotted against
the annealing temperature in Figure 56. As
expected there is a smooth change in the
cation distribution from nearly inverse at
450° C to a more random configuration at
high temperature.

The spectra of MgFe204 in an applied
field are similar to that reported by Sawatsky
et al. (1969). Significantly, the absorption
patterns due to Fe3+ in the distinct crystal
sites are well resolved (Fig. 54). The iso-
mer shift values for Fe3+ in tetrahedral
versus octahedral coordination are 0.31-
0.37mm/sec and 0.43-0.49 mm/sec, respec-
tively.

The broadening of the Z?-site absorption
relative to the A -site absorption, noted pre-
viously, can now be similarly explained as
for other inverse spinels (Sawatsky et al.,

Cation Distributions in MgFe204: XRD
and 57 Fe Mossbauer data.

The structural refinements in which
separate isotropic temperature factors were
refined for each cation site and for oxygen
gave the lowest values for RBragg and Rf
although these statistical criteria were only
slightly improved over a model wherein an
average temperature factor for both cation
sites and for oxygen are refined. The values
for the temperature factors deceased in the
order Btet < B0ct <B oxygen, and the mean
values are Btet = 0.304 ± 0.029 B0ct =
0.360 ± 0.023 and B0Xygen = 0.507
±0.03 1 A2. There is no systematic variation

0.90

- 1%

1 1

1 1 1

S XRD

o Mossbauer

0.80

O^i

n ir\

1 1

o
1 1

i i . T

400

600

800
Temperature,

1000

1200

Fig. 56. A comparison of thermodynamic models
for cation disordering in MgFe204 a) a two term
model with no excess non-configurational en-
tropy of disorder (dashed curve) and b) three term
model, including an excess entropy term (unbro-
ken curve). The XRD data are plotted with one
standard deviation error bars; uncertainty in the
Mossbauer data is 0.008 to 0.0 1 2 and is not shown
for clarity. The three XRD data from 1 100 to
1200°C (open symbols) were not included in the
regressions, because of the possibility of a small
oxygen deficiency in these samples.

GEOPHYSICAL LABORATORY

1969). In the inverse MgFe204, each tetra-
hedral site is surrounded by twelve nearest
Fe3+ neighbors. On the other hand, each
octahedral site is surrounded by only six
tetrahedral Fe3+ neighbors. It is proposed
that the broadening of the B-site absorption
lines is due to the different number of
probable distributions of Fe3+ and Mg on
the six nearest neighbor A -sites. Quite the
opposite effect is proposed for the A-site
Fe3+, because there is no apparent line
broadening at this site. For MgFe204 with
x ~ 0.66, the statistical probabilities that the
B sites have distributions of 6 Fe, 5 Fe 1
Mg, 4 Fe 2Mg and 3 Fe 3Mg in the next
nearest neighbor A sites are 0.24, 0.39,
0.26, and 0.09. For samples that are nearly
inverse with x ~ 1.0, the corresponding
probabilities are 0.78, 0.20, 0.02, 0.0. Thus,
the increased halfwidth of the 5-site with
increasing disorder (cf Fig. 54) is reason-
ably explained in terms of additional hy-
perfine fields due to the next-nearest-neigh-
bor effect.

The values of the inversion parameters
calculated from the area ratios as deter-
mined from the absorption spectra are com-
pared with the values from the Rietveld
refinements in Fig. 57. The mean differ-
ence in the values of x is 0.0056 with a
standard deviation of 0.0102. Thus, the
5 1 Fe Mossbauer data are in excellent agree-
ment with the XRD data, although it is
noted that this agreement would become
less satisfactory if fully ionized atom scat-
tering factors were used in the structural
refinement or if isotropic temperature fac-
tors had not been separately refined for
both cation sites, or if, in fact, only a single
temperature factor was refined.

Thermodynamic model.

The equilibrium values of x in the tem-
perature range 450-1050°C determined
from the XRD Rietveld refinements and
the 57pe Mossbauer data have been fitted
to the expression

-ln# = ccMS-Fe3+ + 2p;t,

(2)

where K is the distribution coefficient
(Navrotsky and Kleppa, 1967). The values
of x were weighted according to their stan-
dard deviation, namely ±0.008 to ±0.012

Comparison of Mossbauer and XRD

0.90

To
a>

o 0.80

0.70

1:1 line

0.70 0.80

x (Rietveld refinements)

0.90

Fig. 57. Comparison of the powder XRD determi-
nations of x from the Rietveld refinements with the
corresponding values determined from the 57Fe
Mossbauer spectra. Data are plotted with one
standard deviation error bars.

for Mossbauer data and ±0.004 for the
XRD data. The results of the fit for this two-
term model, shown plotted in Fig. 56, are
aMg-Fe3+= 26.6 ± 0.4, p = -21.7 ± 0.3 kJ/
mol and X^v = 1.95. There is an improve-
ment in the fit if an additional term repre-
senting a nonconfigurational entropy of
mixing is included. Equation (2) becomes

-In K = aMg-pe3+ - 7oMg-Fe3+ +2p;t , (3)

100

CARNEGIE INSTITUTION

with a Mg-Fe3+ = 16>9 + 2.5 kJ/mol and
aMg-Fe3+= .2.67 ± 1.52 andX^ = 1.10. It
is of interest in the latter case that the values
of x for the heat treated samples at 1100,
1 150 andl250°C are in agreement with the
predicted values using the three-term model.
This latter result is significant, of course, in
terms of whether the high temperature cat-
ion distributions have been quenched at the
respective annealing temperatures.

In the formalism of the thermodynamic
model proposed by O'Neill and Navrotsky
[ 1983, 1984; equation (1)] the nonconfigu-
rational entropy term inferred above is taken
to refer to a vibrational energy contribution
and/or the effect of short-range ordering
(O'Neill and Navrotsky, 1983). In the
latter case it is relevant that significant
positional disorder of Fe3+ and Mg on the
B sub-lattice is inferred from the 57pe
Mossbauer spectra (cf Fig. 54).

In the literature, thermochemical data
required to evaluate the proposed thermo-
dynamic model are sparse. The enthalpy
associated with the change in cation distri-
bution in MgFe204 in the temperature range
700-1200°C has been measured by trans-
posed-temperature-drop calorimetry
(Navrotsky, 1986; Table 3). The experi-
mental value of -5.5 kJ does not agree with
the value of — 1 kJ predicted from the
present study [equations (2), (3)]. How-
ever, it should be noted that the stoichiom-
etry of the MgFe204 used in the thermo-
chemical measurement was not specified.
Further evaluation of the proposed model
for cation disordering in spinels must there-
fore await thermochemical data on well-
characterized samples.

References

Navrotsky, A. and O. J. Kleppa, The thermody-
namics of cation distributions in simple spinels.
/. Inorg. Nucl. Chem., 29, 2701-2714, 1967.

Navrotsky, A., Cation-distribution energetics and
heats of mixing in MgFe204-MgAl204}
2nFe204-2nAl204 and NiAl204-2nAl204
spinels: Study by high- temperature calorim-
etry. Amer. Mineral, 71, 1160-1169, 1986.

Nell J., B. J. Wood, T. O. Mason, High-tempera-
ture cation distributions in Fe304-MgAl204-
MgFe204-FeAl204 spinels from
thermopower and conductivity measurements.
Amer. Mineral, 74, 339-351, 1989.

O'Neill, H. St. C, W. A. Dollase, and C. R. Ross
II, Temperature dependence of the cation dis-
tribution in nickel aluminate (NiAl204) spinel:
a powder XRD study. Submitted to Physics
and Chemistry of Minerals, 1991a.

O'Neill, H St. C, H. Annersten, and D. Virgo D.,
The temperature dependence of the cation
distribution in magnesioferrite (MgFe204)
from powder XRD structural refinements and
Mossbauer spectroscopy. Submitted to Amer.
Mineral, 1991b.

O'Neill, H. St. C. The rates of cation order-
disorder in MgFe204 and MgAl204 spinels.
In preparation, 1991c.

O'Neill, H. St. C, and A. Navrotsky, Simple
spinels: Crystallographic parameters, cation
radii, lattice energies, and cation distribution.
American Mineralogist, 68, 181-194, 1983.

O'Neill, H. St. C, and A. Navrotsky, Cation
distributions and thermodynamic properties
of binary spinel solid soultions. Amer. Min-
eral, 69, 733-753, 1984.

Sawatsky, G.A., F. Van der Wande, and A. H.
Morrish, Mossbauer study of several ferri-
magnetic spinels. Physical Review, 187, 747-
757, 1969.

Virgo, D., S. S. Hafner, Fe^+-Mg order-disorder
in heated orthopyroxenes. Mineral Soc. Amer.
Spec. Paper 2, 67 -87, 1969.

Wiles, D.B., R. A. Young, A new computer pro-
gram for Rietveld analysis of X-ray powder
diffraction patterns. Journal App. Cry stall.,
14, 149-151, 1987.

GEOPHYSICAL LABORATORY

101

Crystallography — Mineral Physics

Predicted High-Pressure Mineral

Structures

with Octahedral Silicon

Robert M. Hazen and Larry W. Finger

Silicates are the most common minerals
on the earth's surface, and they probably
dominate throughout the Earth's mantle.
Many hundreds of silicate structures have
been determined and catalogued (e.g.,
Liebau, 1985), but only about 50 different
structures account for the vast majority of
all crustal silicates (Smyth and Bish, 1988).
A common feature of all these low-pres-
sure mineral structures is the presence of
silicon cations exclusively in 4-coordina-
tion [lYISi by oxide anions.

High-pressure experiments demonstrate
that all common crustal silicates undergo
phase transitions to new structures with 6-
coordinated silicon, [VI]Si, at pressures
between 8 GPa (for pure Si02) to about 30
GPa, which corresponds to the pressure at
the top of the lower mantle. Mineral physi-
cists identify silicon coordination number
as a major crystal chemical difference be-
tween the crust and lower mantle: silicon is
virtually all four coordinated above about
250 km, but is entirely six coordinated
below 670 km. The earth's transition zone,
on the other hand, is marked by the appear-
ance of a group of high-pressure silicates
with both [IVlSi and tVIlSi. The stability of
these minerals is apparently confined to a

rather narrow pressure range from approxi-
mately 10-30 GPa. Within these limits,
however, are silicate structures of remark-
able complexity and great topological in-
terest.

The objectives of this review are to
tabulate all known high-pressure silicate
structures with six coordinated or mixed-
coordinated silicon, to identify crystal
chemical systematics among these struc-
tures, and to predict additional high-pres-
sure silicate structure types.

There are only a dozen known high-
pressure structures with Si06 polyhedra
(Table 13). These silicates can be divided
conveniently into two groups. Above about
25 GPa, corresponding to the Earth's lower
mantle, all silicates studied to date are
observed to transform to one of seven dense
structures, in which all Si is 6-coordinated.
These structures — rutile, perovskite, il-
menite, hollandite, calcium ferrite,
pyrochlore, and K2NiF4 — are well known
room-pressure topologies for transition
metal oxides. In the high-pressure silicate
isomorphs silicon occupies the octahedral
transition metal site, while other cations
may adopt six or greater coordination.

At pressures between about 10 and 20
GPa (in the Earth's transition zone) a sec-
ond group of silicates forms with mixed 4-
and 6-coordination. These phases include
silica-rich modifications of the well known
garnet, pyroxene, and wadeite structures,
as well as complex new magnesium-bear-

102

CARNEGIE INSTITUTION

Table 13. Compositions and calculated densities of high-pressure silicates with Si06 octahedra.

Composition (Mineral Name) Structure Type

p calc*

References

(g/cm3)

A. High-Pressure phases

with Si06

groups only

SiC>2 (Stishovite)

Rutile

4.29

Hill etal. (1983)

CaSi03

Cubic Perovskite

4.25

Mao etal. (1989)

MgSi03

Ortho Perovskite

4.10

Horiuchiertf/. (1987)

MgSi03

Ilmenite

3.81

Horiuchi eet al. (1982)

ZnSi03

Ilmenite

5.25

Ito and Matsui (1974)

KAlSi308

Hollandite

3.91

Yamadaera/. (1984)

BaAl2Si208

Hollandite

5.3

Reid and Ringwood (1969)

CaAl2Si208

Hollandite

3.9

Madon era/. (1989)

NaAlSi04

Calcium Ferrite

3.91

Yamada etal. (1983)

Sc2Si2Ov

Pyrochlore

4.28

Reid etal. (1977)

In22Si207

Pyrochlore

6.34

Reid etal. (1977)

Ca2Si207

K2NiF4

3.56

Liu (1978)

B. High-Pressure phases with SiC>6 +

Si04 groups

MgSi03 (Majorite)

Garnet

3.51

Angela al. (1989)

MnSi03

Garnet

4.32

Fujino etal. (1986)

Na(Mgo.5Sio.5)Si206

Pyroxene

3.28

Angel etal. (1988)

K2Si409

Wadeite

3.09

Swanson and Prewitt (1983)

Mgi4Si5024

Anhydrous Phase B 3.44

Finger etal. (1991)

Mgi2SUOi9(OH)2

Phase B

3.37

Finger etal. (1991)

*p calc = density calculated from unit-cell parameters at room pressure and temperature.

ing phases designated "phase B" and "an-
hydrous phase B." [A third group of room-
pressure [VllSi silicates, detailed by Finger
and Hazen (1991) but not considered here,
includes silicon phosphates with relatively
open framework structures.]

The twelve high-pressure structures
listed in Table 13 display a wide range of
linkages between Si octahedra and other
polyhedra. In stishovite, hollandite, and
pyroxene a combination of edge- and cor-
ner-sharing is observed, but in phase B and
anhydrous phase B each Si octahedron
shares all 12 edges with adjacent Mg octa-
hedra. In pyrochlore, garnet, and wadeite
the Si octahedra form part of a comer-

linked framework, but additional cations in
eight or greater coordination share edges
and faces with the octahedra. Ilmenite pre-
sents yet a different topology, with unusual
face sharing between Mg and Si octahedra,
as well as corner and edge sharing. Despite
the differing polyhedral linkages, the size
and shape of Si06 polyhedra are similar in
all twelve high-pressure compounds. Poly-
hedral volumes at room pressure, for ex-
ample, vary by only about ± 4% from an
average 7.67-A3 value. All Si octahedra are
close to regular (i.e., distortion indices are
small) relative to the range observed for
many divalent and trivalent cation octahe-
dra. The observed tendency of silicon to

GEOPHYSICAL LABORATORY

Table 14. Predicted lVI^Si structures that conform to more than one criterion.

103

Formula

Structure

(MgSi)02(OH)2

Diaspore

Ga^SiOg

Ga4Si7O20

MgioSi30i6

Aerugite

CaSi205

Sphene

MgSi(OH)6

Stottite/Gibbsite

BaSU09

Benitoite

Fe2Si05

Pseudobrookite

* Criteria for predicting tVIlSi structures:

1. Edge-sharing octahedral chains

2. Germanate isomorphs

3. Ti, Mn, and Fe oxides

4. Substitution of (Mg + Si) for 2A1

5. System Mg-Si-O-H

adopt highly regular coordination leads us
to conclude that silicon will not readily
adopt exotic 5- or 7-coordination or highly
distorted 4- or 6-coordination groups at
high pressure in either crystalline or amor-
phous condensed phases.

Five systematic relations among the
structures in Table 1 3 can be used to predict
other possible high-pressure silicates (see
also Table 14). Each of these criteria can be
used to predict other potential [VI] Si phases.

Fig. 58. Relationships among the rutile (A),
IrSe2(B), ramsdellite(C), hollandite (D),
psiomelane (E), and a hypothetical composite
structure (F), after Bursill (1979).

1. Three structure types (rutile,
hollandite, calcium ferrite) are formed from
edge-sharing chains of silicon octahedra.
This relation points to other likely structure
types, all of which incorporate edge-shar-
ing octahedral chains linked to adjacent
strips by corner sharing, as systematized by
Wadsley (1964), Bursill and Hyde (1972),
and Bursill and Hyde (1979). Rutile has
single chains, leading to 1 x 1 square chan-
nels, while hollandite and calcium ferrite
have double chains, yielding larger chan-
nels. Many similar octahedral chain struc-
tures, such as ramsdellite (1x2) and
psilomelane (2 x 3) are also known (Fig.
58), and each of these could provide a
topology suitable for silicon in 6-coordina-
tion.

2. Nine of the twelve known [VllSi
high-pressure structure types were first
synthesized as germinates at lower pres-
sures. A systematic search of the Inorganic
Crystal Structure Data Base (ICSD, FIZ
Karlsruhe distributors) for germanates with
[ vrlGe in systems containing the additional
cations Na, K, Mg, Fe, Ca, Al, Ti, Si, and P

104

CARNEGIE INSTITUTION

revealed 25 structure types, only nine of
which have known silicate analogs (Finger
andHazen, 1991). A number of these com-
pounds, including Fe4Ge209, FesGe30i8,
CaGe205,Ca2Ge70i6,Ca4Ge30io(H20),
and K2BaGegOi8, are good prospects for
high-pressure silicate analogs.

3. All seven high-pressure tVIlSi struc-
tures without tetrahedral Si are isomorphs
of room-pressure oxides with trivalent or
tetravalent transition metals (Ti, Mn, or Fe)
in octahedral coordination. Structures of
other binary oxides with octahedral tita-
nium, manganese, or iron may also repre-
sent possible topologies for mantle miner-
als. Particularly relevant in this context are
the structures of CaSi205 with the sphene
structure, Fe2Si05 or Al2Si05 with the
pseudobrookite structure, and CaSi409 with
the benitoite structure.

4. High-pressure ilmenite, garnet, and
pyroxene forms of magnesium-bearing sili-
cates are all related to room-pressure phases
by the substitution of octahedral Mg and Si
for a pair of aluminum cations. Similar
substitutions might occur at high pressure
in several other common rock-forming
minerals, including kyanite, staurolite,
pseudobrookite, lawsonite, cordierite,
clinozoisite, gibbsite, and diaspore. Note
that this substitution scheme will not work
for many common aluminum-bearing min-
erals with mixed 4- and 6-coordinated alu-
minum. The substitution in muscovite
[K[VI]Al2[IV](AlSi3)Oio(OH2)], for ex-
ample, would yield the magnesian mica
celadonite,K[VI](MgAl)[lV]Si4Oio(OH)2,
in which all Si is tetrahedrally coordinated.
Octahedral Al, thus, must constitute more

than two-thirds of all aluminum to produce
a [^1] Si phase by the substitution 2A1 — >
(Mg + Si).

5 . Finger and Pre witt ( 1 990) documented
the close structural relations among a num-
ber of hydrous and anhydrous magnesium
silicates, and used those systematics to
propose several as yet unobserved struc-
tures, including high-pressure hydrous
phases with octahedral silicon. They rec-
ognized that several known phases, includ-
ing chondrodite, humite, forsterite, phase
B, and anhydrous phase B, are members of
a large group of homologous magnesium
silicates that can be represented by the
general formula:

m[Mg4«+2[][VJSi2«08«(OH)4]Mg6/2+4-
2mod(n,2)[YIlSin+mod(n,2)OSn+4

where mod{n,2) is the remainder when n is
divided by 2. Finger and Prewitt (1990)
examined cases where n = 1 ,2,3,4, °° and m
= 1,2, oo. Structures with octahedral silicon
result for all cases where m is not infinity.
Of special interest is the proposed structure
of "superhydrous phase B," a compound
predicted by the logical progression from
Mgi4Si5024 (anhydrous phase B) to
Mgi2Si40i9(OH)2 (phase B) to
MgioSi30i4(OH)4. Gasparik (1990) sug-
gested that an as yet unanalyzed hydrous
magnesium silicate synthesized at 1 8.6 GPa
and 1600 °C possesses this structure, and
further studies on that material are in
progress.

Several structure types appear to follow
two of the five very different systematic
trends. These structures, therefore, de-

GEOPHYSICAL LABORATORY

105

mand further study. Of special interest to
earth scientists are CaSi20s with the titanite
structure, Fe2SiOs with the pseudobrookite
structure, and Mg ioSi30 16 with the aerugite
structure. Each of these phases, or their
isomorphs with other cations replacing Ca,
Mg, and Fe, might be represented in the
Earth's mantle. In fact, Stebbins and
Kanzaki (1991) mention the existence of
titanite-type CaSi20s in some of their run
products, though identification of this phase
was provisional.

Also worthy of further study are the
proposed hydrous phases MgSi02(OH)2
and MgSi(OH)6, which are isomorphs of
diaspore and stottite, respectively. Such li-
nen phases would be expected to occur
only locally in the earth's deep interior, but
their presence, integrated over the Earth's
volume, could represent a major respository
of water.

Most common rock-forming cations,
including Na, Mg, Fe, Ca, Mn, Al, Ti, and
Si, are small enough to fit into the tetrahe-
dral or octahedral interstices of a close-
packed oxygen net. However, the presence
of many other cations, including H, B, K,
Rb, Pb, rare earths, and U, could disrupt the
close-packed array and lead to other, as yet
unrecognized, structure types. The gal-
lium and barium silicates, Ga4SiOs,
Ga4SiyO20, and BaSi409, are just three of
the dozens of possible new [VlJSi struc-
tures likely to be observed as high-pressure
investigations extend beyond the traditional
rock-forming elements. These structures
are not likely to play a significant role in
mantle mineralogy, but they will provide a
more complete understanding of the crys-
tal chemistry of octahedral silicon.

Conclusions

Is the earth's deep interior
mineralogically simple? Are there only a
few dominant structure types, or is there an
unrecognized complexity in the crystal
chemistry of octahedral silicon? There are
hundreds of different crustal silicates with
PYlSi, but only a dozen high-pressure [VI]Si
structures have been produced. This dis-
parity may reflect the relatively small num-
ber of high-pressure studies, but it also
arises, at least in part, from the nature of
oxygen packing. Numerous crustal sili-
cates, from the commonest minerals quartz
and feldspar to the dozens of zeolites and
other framework silicates, possess open,
low-density topologies with correspond-
ingly loose packing of oxygen. There are
no obvious limits to the variety of silicates
based on irregular oxygen packing.

Volume constraints imposed by high
pressure, however, favor structures with
approximately close-packed oxygens.
These restrictions on anion topology re-
duce the number of possible cation con-
figurations as well, and it is thus antici-
pated that the number of different struc-
tural topologies in the earth's deep interior
will be much smaller than at the surface.
Dense, close-packed, and for the most part
high-symmetry structures, such as those
represented by the seven known topologies
with all tvrlSi, will predominate. Never-
theless, within these restrictions there ex-
ists opportunity for considerable structural
diversity owing to three factors - reversible
phase transitions, cation positional order-
ing, and modularity, particularly based on

106

CARNEGIE INSTITUTION

different close-packed layer stacking se-
quences. This potential diversity is only
hinted at by the known phases.

Several of the known high-pressure
types, including perovskite, K2NiF4, and
pyrochlore, can adopt numerous structural
variants based on slight changes in lattice
distortions and cation distribution. The
perovskite structure, in particular, can un-
dergo dozens of phase transitions based on
octahedral tilting, cation ordering, cation
displacements, and anion defects (Megaw,
1973; Hazen, 1988). We must study pro-
posed mantle phases at the appropriate
conditions of pressure and temperature to
document the equilibrium structural varia-
tions.

Close packing of oxygen leads to modu-
lar structures, with certain features (e.g.,
edge-sharing octahedral chains of rutile;
the double chains of hollandite; the comer-
sharing octahedral sheets of perovskite; the
face-sharing topology of ilmenite) that can
link together in many ways to form ordered
superstructures of great complexity. Such
complexity was recognized by Wadsley
( 1 964) and Bursill and Hyde ( 1 979) in their
descriptions of modular rutile-hollandite-
Ga203 structures, and it is realized in the
homologous series including phase B, an-
hydrous phase B, and several other struc-
tures. Phase B, for example, is based on
oxygen close packing, yet it has 40 inde-
pendent atoms in its asymmetric unit to
yield one of the most complex ternary
silicates yet described. Variations on the
phase B structure could be based on chang-
ing the relative number and position of the
two different structural layers, by introduc-

ing other types of layers, or by staggering
layers to produce clino- and ortho-type
structures as observed in other close -packed
systems, for example, in the biopyriboles
as describedbyThompson(1978) and Smith
(1 982). The structure could be further com-
plicated by element ordering among the 17
different cation sites as Al, Fe Ti, Mn, and
other elements enter the structure in a natu-
ral environment .

The study of octahedrally-coordinated
silicon is still in its infancy, yet clear trends
are beginning to emerge from the scattered
data on diverse structures and composi-
tions. It is now evident that while silicate
perovskite may be the predominant phase
in the Earth's lower mantle, a number of
other dense silicate phases will compete for
elements such as K, Ba, Ca, and Al. It
appears that the earth's transition zone will
display the varied mineralogy of mixed
tVIlSi and PV] Si silicates, including some
of the most complex structures known in
the mineral kingdom. And it is certain that
a detailed understanding of the mantle must
await studies of these fascinating phases at
temperatures and pressures appropriate to
the Earth's dynamic interior.

References

Angel, R. J., T. Gasparik, N. L. Ross, L. W. Finger,
C. T. Prewitt, and R. M. Hazen, A silica-rich
sodium pyroxene phase with six-coordinated
silicon, Nature, 335, 156-158, 1988.

Angel, R. J., L. W. Finger, R. M. Hazen, M.
Kanzaki, D. J. Weidner, and R. C. Liebermann,
Structure and twinning of single-crystal
MgSiC>3 garnet synthesized at 17 GPa and
1800°C, Amer. Mineral, 74, 509-512, 1989.

Bursill, L.A. and B. G. Hyde, Structural relation-
ships between P-gallia, rutile, hollandite,

GEOPHYSICAL LABORATORY

107

psilomelane, ramsdellite, and gallium titanite
type structures, Acta Cryst., B35, 530-538,
1979.
Bursill, L. A., and B. G. Hyde, Rotation faults in
crystals, Nature (London) Phys. Sci., 240, 122-
124, 1972.
Finger, L. W., and R. M. Hazen, Crystal chemistry
of six-coordinated silicon: A key to under-
standing the Earth ' s deep interior, Acta Cryst.,
A47, 1991, in press.
Finger, L.W. and C. T. Prewitt, Predicted compo-
sitions for high-density hydrous magnesium
silicates, Geophys. Res. Lett., 16, 1395-1397,
1990.
Finger, L. W., R. M. Hazen, and C. T. Prewitt,
Crystal structures of Mgi2Si40i9(OH)2 (Phase
B) and Mgi4Si5024 (Phase AnhB), Amer.
Mineral, 76, 1-7,1991.
Fujino, K., H. Momoi, H. Sawamoto, and M.
Kumazawa, Crystal structure and chemistry of
MnSi03 tetragonal garnet, Amer. Mineral., '
71, 781-785, 1986.
Gasparik, T. Phase relations in the transition zone,

/. Geophys Res., 95, 15751-15769, 1990.
Hazen, R. M., Perovskites, Sci. Amer., 74-8 1 , June

1988.
Hill R.J., M. D. Newton, and G. V. Gibbs, A
crystal chemical study of stishovite, /. Sol.
State Chem., 47, 185-200, 1983.
Horiuchi, H., M. Hirano, E. Ito, and Y. Matsui,
MgSiC>3 (ilmenite-type): Single crystal x-ray
diffraction study, Amer. Mineral, 67, 788-793,
1982.
Horiuchi, H., I. Eiji, andD. J. Weidner, Perovskite-
type MgSiC>3: Single-crystal x-ray diffraction
study, Amer. Mineral, 72, 357-360, 1987.
Ito, E., and Y. Matsui, High-pressure synthesis of
ZnSiC>3 ilmenite, Phys. Earth Planet. Int., 9,
344-352, 1974.
Liebau, F., Structural Chemistry of Silicates,

Springer, New York, 1985.
Liu, L., High pressure Ca2SiC>4, the silicate
K2NiF4-isotype with crystalchemical and geo-
physical implications, Phys. Chem. Minerals,
3, 291-299, 1978.
Madon, M, J. Castex, and J. Peyronneau, A new
aluminocalcic high-pressure phase as a pos-
sible host of calcium and aluminum in the
lower mantle, Nature, 342, 422-425, 1989.
Mao, H.K., L. C. Chen, R. J. Hemley, A. P.
Jephcoat, Y. Wu, and W. A. Bassett, Stability
and equation of state of CaSi03-perovskite to
134 GPdiJ.Geophys.Res.,94, 17,889-17,894,
1989.
Megaw, H. D., Crystal Structures: A Working
Approach, Philadelphia, W. B. Saunders, 1973.
Reid, A. F., and A. E. Ringwood, Six-coordinate
silicon: High pressure strontium and barium
aluminosilicates with the hollandite structure,
J. Solid State Chem. 1, 6-9, 1969.

Reid, A. F., C. Li, and A. E. Ringwood, High-
pressure silicate pyrochlores, SC2S12O7 and
In2Si207, Solid State Chem., 20, 219-226,
1977.

Smith, J. V., Geometrical and Structural Crystal-
lography, Wiley, New York, 1982.

Smyth, J. R., and D. L. Bish, Crystal Structures
and Cation Sites of the Rock-Forming Miner-
als, Allen and Unwin, Winchester, Massachu-
setts, 1988.

Stebbins, J. F., and M. Kanzaki, Local structure
and chemical shifts for six-coordinated silicon
in high-pressure mantle phases, Science, 251,
294-298, 1991.

Swanson, D.K., and C. T. Prewitt, The crystal
structure of K2SiVISi3IVC>9, Amer. Mineral,
65,581-585,1983.

Thompson, J. B., Biopyriboles and polysomatic
series, Amer. Mineral, 63, 239-249, 1978.

Wadsley, A. D.,Non-Stoichiometric Compounds,
L. Mandelcoin, ed., p. Ill, New York, Aca-
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Yamada, H., Y. Matsui, andE. Ito, Crystal-chemi-
cal characterization of NaAlSi04 with the
CaFe204 structure, Mineral. Mag., 47, 177-
181, 1983.

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29-34, 1984.

SlMULTANOUS HlGH P-^

Diffraction Measurements

OF (Fe,Mg)Si03-PEROVSKITE

and (Fe,Mg)0 Magnesiowustite:

Implications for Lower

Mantle Composition

Yingwei Fei, Ho-Kwang Mao, Russell J.
Hemley, and Jinfu Shu

(Fe,Mg)Si03-perovskite and (Fe,Mg)0
magnesiowustite are likely stable phases in
the Earth's lower mantle. The thermal prop-
erties of those phases are of critical impor-
tance for constraining the composition of
the lower mantle. In this paper, we report
simultaneous high P-T synchrotron x-ray
diffraction measurements of (Fe,Mg)Si03

108

CARNEGIE INSTITUTION

>n Mw

C/)

Mw200

Mw220

Au200

Au111

'•'. Au311

'• Au220 Mw311 :".*

19.8 GPa, 310 K

Energy, keV

Fig. 59. Representative energy-dispersive x-ray diffraction spectra (20 = 15°) of magnesiowustite
(Mg.6Fe.4)0 at three different P-T conditions.

perovskite and (Fe,Mg)0 magnesiowustite.
These data provided the first direct mea-
surements of the effect of pressure on the
thermal expansivity of these minerals. In
the discussion we also examine the impli-
cations of these results for the composition
and mineralogy of the lower mantle.

Experimental Methods

The samples used in the experiments
are (Feo.4Mgo.6)0 magnesiowustite, syn-
thesized from mixtures of Fe2C>3 and MgO

at a temperature of 1573 K and an oxygen
fugacity (fen) of lO10-8 bar (Rosenhauer et
a/., 1976), and (Feo.iMgo.9)Si03 perovskite,
synthesized from synthetic pyroxene by
laser heating at 40 GPa in diamond-anvil
cell (Mao etal, 1991).

The experiments were carried out using
a high-temperature diamond-anvil cell,
made of inconel, and synchrotron radia-
tion. A nickel alloy (Rene 41) gasket with
thickness of 200 Jim was preindented to a
pressure of 17 GPa. The powder sample
was placed in the 250-|im-diameter sample
chamber, filling less than one-third of the

GEOPHYSICAL LABORATORY

109

volume. Gold foil and ruby grains were
placed in the sample chamber as pressure
calibrants at high temperature (Anderson
et al. , 1 989) and at room temperature (Mao
etal. , 1 986), respectively. The sample cham-
ber was then filled with neon gas at 200
MPa in a high-pressure gas-loading device,
and sealed at a pressure of 2 GPa. The neon
served as a quasi-hydrostatic pressure trans-
mitting medium over the pressure range of
measurement.

The sample was heated with an external
platinum-wire resistance heater (Mao et
al, 1991). The heater was placed on the
cylinder of the cell. Temperatures were
measured with a chromel-alumel thermo-
couple, while pressures were determined
by measuring the lattice parameter of gold.
During the experiment, pressure usually
decreases with increasing temperature at
the rate of about 5 GPa/100 K.

Polychromatic (white) wiggler synchro-
tron x-radiation at the National Synchro-
tron Light Source, Brookhaven National
Laboratory was used for the energy-disper-
sive x-ray diffraction measurements. The
diffraction data were collected with an in-
trinsic germanium solid-state detector at a
fixed 20 angle of 15°(± 0.005°). The energy
was calibrated by using known energies of
x-ray emission lines (Ka and Kp) of Mn,
Cu, Rb, Mo, Ag, Ba, and Tb. The 20 angle
was calibrated by collecting the diffraction
pattern of platinum. The experimental con-
ditions were optimized by considering the
x-ray beam size, data collecting time, and
slit size for the detector. With a 60-|im
beam spot, a complete diffraction pattern
of magnesiowustite with reasonable peak
counts can be obtained in about five min-

utes. Figure 59 shows three typical spectra
of magnesiowustite with internal standard
gold collected at different P-T conditions.
All diffraction patterns show at least four
sharp diffraction lines, 111, 200, 220, and
311, of magnesiowustite. For theperovskite,
the diffraction was carried out using mono-
chromatic synchrotron x-ray and film meth-
ods (Mao et al., 1991). These techniques
were used because of the need for high
resolution to resolve multiplets in the dif-
fraction patterns arising from the or-
thorhombic distortion of the perovskite.

Results and Discussions

Lattice parameters of the
magnesiowustite were determined from
diffraction lines 111, 200, 220, and 311 by
using a peak-fitting program. For
perovskite, diffraction peaks (mainly 020,
112, and 200) were measured by both
manual and computerized film-reading
methods. The experimental results are plot-
ted in Figures 60 and 61 . The uncertainties
in pressure result from the measurements
of lattice parameter of gold which was used
as an internal high-pressure calibrant at
high temperature. The error of ± 0.0015 A
in the lattice parameter of gold, which is the
typical measurement uncertainty in the ex-
periments, corresponds to about ± 0.30
GPa in pressure.

To construct a P-V-T equation of state
from a combination of 1-bar thermal ex-
pansion data, 300-K compression data, and
simultaneous high P-T volume data, we
express pressure in a general form

/W) = /W300K) + />*, (1)

110

CARNEGIE INSTITUTION

o
o

CO
CL

P
>

10 15

Pressure, GPa

Fig. 60. Volume at high P-T relative to the 300 K
isotherm. Experimental data are from Mao et al
[1991] (crosses); Knittle et al, [1986] (solid
squares); and Wang et al, [1991] (open squares).
The lines are calculated isotherms.

o
o

a."
>

2.5-1

2.0_

1.5-

1.0-

0.0-

S93"*-

600

0, „,, 4- _^«76

565

5*3

*"1 «•

323 -}-31Q

500K

300K

—T
25

- 1-
30

Pressure, GPa

Fig. 61. Volume at high P-T relative to the 300 K
isotherm. Experimental data are from Fei et al
[1991b] (crosses); and Suzuki [1975] (solid
squares). The lines are calculated isotherms.

where P(V,300K) can be expressed by the
standard third-order Birch-Murnaghan form

P = h

4#-(#I-^-^fe

(2)

where Kto and KTq are the isothermal bulk
modulus and its pressure derivative at room
temperature, respectively. The thermal pres-
sure can be modeled in two ways as dis-
cussed below.

By using the thermodynamic identity,
the thermal pressure can be calculated by
integrating the oKv i.e.,

/>th =

/300K

aKjiXT

(3)

where a is the thermal expansivity and KT
is isothermal bulk modulus at T and V of
interest. To parameterize the thermal pres-
sure, we start with an assumption that (BK^
dT)p is independent of temperature, and
obtain (dKflT), = -2 .7 '(±0.3) x 102 GPa/K
for magnesiowiistite and (dK^dT^p = -
6.3(±0.5) x 10 2 GPa/K for perovskite by
fitting our P-V-T data to the thermal pres-
sure model. A detailed discussion of these
results is given in Mao et al ( 1 99 1 ) and Fei
etal (1991b).

The Anderson-Griineisen parameter ST
is commonly used to measure the change in
thermal expansivity at high P-T. It is de-
fined by (Anderson, 1967)

<5r =

3lna

_ 1

lMT\

[dlnV It aKA dT

(4)

The parameters, 5r for magnesiowiistite
and perovskite derived from our data are
listed in Table 15. They decrease with in-
creasing temperature below the Debye tem-
perature and approach a constant value at
high temperature. The ST values for
magnesiowiistite and perovskite are 4.3
and 6.5, respectively, above the Debye
temperatures.

GEOPHYSICAL LABORATORY

111

Table 15. Temperature variation of some thermodynamic parameters for magnesiowiistite and
perovskite

Perovskite

Magnesiowiistite

T,K

a(106)

KT, GPa

ocKt

a(106) tfr,GPa

aKj

300

21.90

260.9

57.15

11.02

7.48

31.32

157.0

49.16

5.49

1.97

400

28.94

254.6

73.68

8.55

5.28

35.87

154.3

55.33

4.88

1.57

500

33.00

248.3

81.94

7.69

4.50

38.47

151.6

58.31

4.63

1.40

600

35.90

242.0

86.86

7.25

4.10

40.31

148.9

60.00

4.50

1.30

700

38.24

235.7

90.11

6.99

3.86

41.79

146.2

61.07

4.42

1.24

800

40.28

229.4

92.40

6.82

3.70

43.07

143.5

61.79

4.37

1.19

900

42.15

223.1

94.04

6.70

3.58

44.25

140.8

62.28

4.34

1.15

1000

43.92

216.8

95.21

6.62

3.50

45.35

138.1

62.60

4.31

1.13

1100

45.62

210.5

96.02

6.56

3.45

46.41

135.4

62.81

4.30

1.11

1200

47.27

204.2

96.52

6.53

3.42

47.43

132.7

62.92

4.29

1.09

1300

48.89

197.9

96.74

6.51

3.41

48.44

130.0

62.94

4.29

1.09

1400

50.48

191.6

96.71

6.51

3.41

49.43

127.3

62.90

4.29

1.08

1500

52.05

185.3

96.45

6.53

3.43

50.40

124.6

62.78

4.30

1.08

1600

53.62

179.0

95.96

6.57

3.47

51.37

121.9

62.60

4.31

1.09

1700

55.17

172.7

95.26

6.61

3.52

52.33

119.2

62.36

4.33

1.10

The thermal contribution can also be
calculated by the Mie-Gruneisen relation,

^h = Qect, eD) - £(3ook, eD)]

where E(T,QD) is the harmonic internal en-
ergy calculated from either a Debye model
or a single Einstein oscillator model
(Zharkov and Kalinin, 1971), which are
equivalent in the high-temperature limit.
The Debye temperature 6D and Griineisen
parameter /are considered to be functions
of volume only: y= -dkiOJdlnV , with the
volume dependence of y given by q =9lny
/3lnK The model parameters, 0D, y, and q,
can be obtained by fitting the experimental
P-V-T data to the Mie-Griineisen equation
of state.

There is some uncertainty in the Debye
temperature Qm for perovskite. As a result
of the non-linear character of the fit, the
parameters are correlated and there is a
trade-off among the best-fit values. Figure
61 illustrates the trade-off between y0 and q
for values of Bm ranging from 725 to 1025

K for the silicate perovskite. The circles
indicate the best fit to the experimental data
when both y0 and q are simultaneously
optimized. Notably, the q of 3 . 3 (at y0 = 1 .70
and 6m = 725 K) for perovskite obtained in
this analysis is considerably higher than
many other materials {e.g., q = 0-1 is com-
monly found in shock- wave studies). How-
ever, the high q value is consistent with
high <57-value derived from an independent
method of analysis (Mao et al. , 1 99 1 ) using
the thermodynamic relation

q = 8T + 1

(6)

when (dlnCydlnVOy, = 0, which is valid at
high temperature.

The trade-off between the values of K.

and Kn' that fit the static compression data
has been examined previously (Mao et al.,
1 99 1 ). That work showed that the assump-
tion that Kn* = 4 yields Kw = 26 1 GPa. If a
lower value for the bulk modulus is as-
sumed, a higher value for K^ is required to
be consistent with the static compression

112

CARNEGIE INSTITUTION

data. For example, adopting Kw = 247
GPa, as obtained from Brillouin scattering
measurements at zero pressure by Yeganeh-
Haeri et al. (1989), requires Kn* = 5.5.
Because of the identity relating K^ and q
[equation (6)], it is useful to consider the
effect of a higher assumed value for A^' on
the thermal properties. A fit to the experi-
mental data with the assumption of Kn' =
5.5 yields q = 1. 8 at y0 =1.70, with 0D0 = 725
K (Fig. 62). The difference in densities
calculated for perovskite at high P and T
with Kw' varied over this range increases
with pressure: for example, at 60 GPa K^
= 4 gives 0.8% higher density than that
calculated with K^ = 5.5.

The parameter values for calculating
the P-V-T relations of perovskite and
magnesiowiistite are summarized in Table
16. Isotherms calculated with the Mie-
Griineisen relation in the range of the x-ray
diffraction measurements are shown in Fig-
ures 62 and 63. For magnesiowiistite, both
equations (3) and (5) both predict consis-
tent P-V-T relations up to 140 GPa and
3000 K(Fei etaL, 1991b). For perovskite,
however, equation (3) predicts higher vol-
ume at high P-T than equation (5) (about
0.5% higher at 60 GPa and 2000 K). These
differences can result in large uncertainties
in the lower mantle composition.

Fig. 62. The trade-off between y0 and q at various
given Debye temperature 0D0. The circles indi-
cate the best fit to the experimental data when
both y0 and q are simultaneously optimized. The
solid lines are obtained when Kw = 261 GPa and
AYo = 4 are used. The dashed lines are obtained
when K-ro - 247 GPa and KTq = 5.5 are used. (See
text for discussion.)

A detailed comparison of the densities
calculated from P-V-T equations of state
presented above and that determined by
seismology (e.g., PREM, Dziewonski and
Anderson, 1981) indicates that a large
perovskite component relative to
magnesiowusite with an upper mantle Fe/
Mg ratio (Ring wood, 1 975) matches PREM
to within 1 % throughout the entire lower

Table 16. Parameters of the thermal equations of state of perovskite and magnesiowiistite

Parameters

(Fe,Mg)SiC>3-perovskite

(Fe,Mg)0-magnesiowustite

Vo, cm3/mol

24A6+0MXFe

11.25+1.02XFe

Km GPa

261(4)

160-7.5XFe

Kto

(dK^dT)P, GPa/K

-6.3(5) x lO-2

-2.7(5) x lO"2

(dKiftnv, GPa/K

-2.5(3) x 10-2

-0.2(2) x lO"2

6.5(5)

4.3(5)

Oqo, K

725(25)

500(20)

1.70(5)

1.50(5)

3.3(5)

1.1(5)

GEOPHYSICAL LABORATORY

113

mantle (see Hemley et al., 1991). This
implies a silica enrichment in the lower
relative to the upper mantle. The conclu-
sion is dependent on the geotherms and the
thermoelastic parameters. Figure 63 shows
the Fe/(Mg+Fe) and Si/Mg ratios for a best
fit of density between the mineralogical
model and PREM in the top portion of the
lower mande (pressure range of 24-60 GPa).
The composition trade-off between tem-
perature and density is demonstrated by
varying the adiabatic temperature at 670
km from 1 800 K to 2000 K and its gradient
from 7 K/GPa to 10 K/GPa. Abest fit of X

= 0.1 andXs.= 1.0 is obtained for the 2000-
K adiabat. The Xsi value is not as well
constrained by the density analysis. How-
ever, the bulk modulus, the slope of density
profile divided by density, provides a con-
straint on Si/Mg. The minimum of XSi is
always close to 1 (i.e., pure perovskite
composition) for the equation of state of
perovskite used in this study. The mini-
mum shifts to more olivine-enriched
composition if the volume dependence of
the thermal expansivity of perovskite de-
creases.

Recently, we also showed that the cal-
culated phase relations in the MgO-FeO-
Si02 system under lower mantle condi-
tions are sensitive to the volume depen-
dence of the thermal expansion coefficient
of component phases, with the value for the
(Fe,Mg)Si03 perovskite being especially
significant (Fei and Hemley, 1991). We
find that <5rof at least 6 (q of 2.9) is required
to match the observed phase boundary for
Mg2Si04 (spinel) = MgSi03 (perovskite)
+ MgO (periclase). A high value for 5j is
also found to be consistent with the experi-
mental phase equilibrium data (Ito and
Takahashi, 1989) and element partitioning

0.20

0.15

0.10

0.05 -

decreasing q

q=2.3
18O0K

10K/GPa 9=3.3
1800K

(7=3.3
2000K
10K/GPa

g=3.3
10K/GPa 1800K

7K/GPa

0.00

0.5

0.6

0.7 0.8

Xsi

0.9 1 .0

Fig. 63. Misfits between the mineralogical model
and seismic data for density as functions of Xpe and
Xsi. The Fe/Mg and Si/Mg ratios for a best fit of
density are calculated by varying the adiabatic
temperature at 670 km from 1 800 K to 2000 K and
its gradient from 7 K/GPa to 10 K/GPa. The effect
of the volume dependence q for perovskite is also
indicated by the arrow.

data (Fei et al., 1991a) in the system
Mg2Si04-Fe2Si04. The stability field of
perovskite shifts to higher pressure and the
Mg-rich region with decreasing &r value.
A lower value of &r would give rise to two
complete solid solutions between spinel
and magnesiowustite and between
perovskite and magnesiowustite, in contra-
diction to experimental results in this sys-
tem. Increasing XFe and temperature, and
decreasing <5r, are shown to decrease sig-
nificantly the stability field of perovskite.
The primary conclusions of this study
concern the top of the lower mantle (from
670 - 1000 km depth). Comparison be-
tween the extrapolated equation of state
data and seismic results at greater depths
indicates that the compositional arguments
put forth here do apply to the bottom of the
lower mantle; that is, the conclusion that a
lower mantle assemblage dominated by

114

CARNEGIE INSTITUTION

perovskite with an^Fe of -0.1 is consistent
with the seismic data. The main uncertain-
ties in this lower mantle composition model
may rise from both temperature and pres-
sure extrapolations in the thermal pressure
models. The two models for perovskite
used in the analyses result in 0.5% differ-
ence in desity which accounts for about 2%
difference in iron content by fitting to the
PREM densities. Uncertainties in Kj can
result in more significant changes in the
composition model, especially to the Si/
Mg ratio argument. More accurate models
of the composition and mineralogy in the
deeper portions of the mantle would re-
quire measurements at higher P-T condi-
tions.

References

Anderson, O. L., Equation for thermal expansivity
in planetary interiors, J. Geophys. Res., 72,
3661, 1967.

Anderson, O. L., D. G. Isaak, and S. Yamamoto,
Anharmonicity and the equation of state for
gold, /. Appl. Phys., 65, 1534-1543, 1989.

Dziewonski, A. M., and D. L. Anderson, Prelimi-
nary reference earth model, Phys . Earth Planet.
Interiors, 25, 297-356,1981.

Fei, Y., andR. J. Hemley, Stability of (Fe,Mg)Si03-
perovskite in the lower mantle, Geophys. Res.
Lett., in press, 1991.

Fei, Y., H. K. Mao, and B. O. Mysen, Experimen-
tal determination of element partitioning and
calculation of phase relations in the MgO-
FeO-Si02 system at high pressure and high
temperature, /. Geophys. Res., 96, 2157-
2169, 1991a.

Fei, Y., H. K. Mao, J. Shu, and J. Hu, P-V-T
equation of state of magnesiowustite
(Mgo6Feo4)0, Phys. Chem. Miner., submit-
ted, 1991b.

Hemley, R. J., Y. Fei, andH. K. Mao, Constraints
on lowermantle composition from P-V-T mta-
surements of (Fe,Mg)SiC>3-perovskite and
(Fe,Mg)0, High Pressure Research in Min-
eral Physics: Application to Earth and Plan-
etary Science, edited by Y. Syono and M. H.
Manghnani eds., submitted, 1991.

Ito, E., and E. Takahashi, Postspinel transforma-
tions in the system Mg2Si04-Fe2Si04 and
some geophysical implications, /. Geophys.
Res., 94, 10,637-10,646, 1989.

Knittle, E., R. Jeanloz, and G. L. Smith, The
thermal expansion of silicate perovskite and
stratification of the Earth's mantle, Nature,
319, 214-216, 1986.

Mao, H. K., J. Xu, and P. M. Bell, Calibration of
the ruby pressure gauge to 800 kbar under
quasihydrostatic conditions,/. Geophys. Res.,
97,4673-4676,1986.

Mao, H. K., R. J. Hemley, Y. Fei, J. F. Shu, L. C.
Chen, A. P. Jephcoat, Y. Wu, and W. A.
Bassett, Effect of pressure, temperature, and
composition on lattice parameters and density
of (Fe,Mg)SiC>3-perovskites to 30 GPa, J.
Geophys. Res., in press, 1991.

Ring wood, A. E., Composition and Petrology of
the Earth's Mantle, pp. 618, McGraw-Hill,
New York, 1975.

Rosenhauer, M., H. K. Mao, and E. Woermann,
Compressibility of magnesiowustite
(Feo 4Mgo 6)0 to 264 kbar, Carnegie Inst.
Washington Year Book , 75, 513-515, 1976.

Suzuki, I., Thermal expansion of periclase and
olivine and their anharmonic properties, J.
Phys. Earth, 23, 145-159, 1975.

Wang, Y., D. J. Weidner, R. C. Liebermann, X.
Liu, J. Ko, M. T. Vaughan, Y. Zhao, A.
Yeganeh-Haeri, and R. E. G. Pacalo, Phase
transition and thermal expansion of MgSi03
perovskite, Science, 251, 410-413, 1991.

Yeganeh-Haeri, A., D. J. Weidner, and E. Ito,
Elasticity of MgSiC>3 in the perovskite struc-
ture, Science, 243, 787-789, 1989.

Zharkov, V. N., and V. A. Kalinin, Equation of
State for Solids at High Pressures and
Temperatures, pp. 257, Consultants Bureau,
New York, 1971.

GEOPHYSICAL LABORATORY

115

High-Pressure Crystal Chemistry

of Iron-Free wadsleyite,

P-MG2S1O4

Jinmin Zhang, Robert M. Hazen, and
Jaidong Ko*

Wadsleyite, P-(Mg,Fe)2Si04, may be
the most abundant mineral in the upper
mantle between 400 and 550 km, and the
phase transformation of olivine to
wadsleyite may explain the 400-km dis-
continuity. Since this phase was first found
to occur at high pressure, much work has
been done to relate its crystal structure and
properties to geophysically important prob-
lems. For example, the 01 and even 02
positions have been considered to be pos-
sible sites for protonation; therefore, the
wadsleyite phase may be a repository of
water in the mantle (Smith, 1987; Downs,
1989).

The wadsleyite structure has three crys-
tallographically non-equivalent octahedral
sites, M7, Ml and M3. Wadsleyite in the
mantle is believed to contain -10% Fe2+,
which partitions in the order M3>M1>M2
(Finger et al., in preparation). This prefer-
ence should depend on the volume, the
distortion and crystal field stabilization
energy (CFSE) of each site. The character-
istics of these cation sites in an iron-free
phase at high pressure are important to
understanding the distribution of Fe2+ in
wadsleyite in the mantle. In this paper, we
report the result of structure refinements of
P-Mg2Si04 at five pressures up to 4.84

* Department of Earth and Space Science, State
University of New York at Stony Brook

GPa, with particular attention to the prop-
erties of the octahedral sites. This study,
together with work in progress on Fe-con-
taining wadsleyites, will help to define the
properties of this phase in the mantle.

The single crystals used in this study
were synthesized by Jaidong Ko at the
High-Pressure Laboratory of the State Uni-
versity of New York at Stony Brook. The
sample was produced at a pressure of about
16 GPa and a temperature of 1400°C.

A subequant crystal 0.06 x 0.12 x 0.13
mm in size was selected, and wasexamined
optically and by x-ray diffraction at room
conditions. Three intensity data sets were
collected and several unit-cell parameter
measurements were made on this crystal
before it was crushed when trying to in-
crease pressure after the run at 2.88 GPa.
Another crystal 0.04 x 0.06 x 0.06 mm in
size was used for higher pressure measure-
ments.

The selected crystal and a small piece of
fluorite crystal were mounted in a dia-
mond-anvil cell designed for single -crystal
x-ray diffraction studies (Hazen and Fin-
ger, 1982). The fluorite crystal was used as
an internal standard of pressure, following
the method of Hazen and Finger ( 1 98 1 ). An
Inconel 750X gasket with 0.40 mm diam-
eter hole was centered over one 0.60-mm
diamond anvil; the crystal was then affixed
to the anvil face inside the hole with a small
dot of the alcohol insoluble fraction of
vaseline petroleum jelly. A mixture of 4:1
methanol rethanol was used as the hydro-
static pressure medium.

All x-ray measurements except for the 0
and 1.16 GPa intensity data collections
were performed with a Picker automated

116

CARNEGIE INSTITUTION

Table 17. Crystallographic data for iron-free wadsleyite

Pressure, GPa

Paramter

1.16

1.81

2.88

4.84

crystal size, mm

0.06x0.12x0.13

0.04x0.06x0.06

IH> cm-1

11.16

11.25

11.29

11.35

11.48

Range of TO)

0.92-0.94

0.92-0.94 0.93-0.94

0.93-0.94

0.95-0.96

RinP

0.060

0.069

0.057

0.078

Number of data

307

294

303

286

239

R J3) all data

0.056

0.055

0.056

0.059

0.060

/?(4) all data

0.089

0.094

0.082

0.072

0.114

Number F0>2gf

236

216

244

242

157

0.055

0.054

0.055

0.057

0.058

0.068

0.069

0.061

0.057

0.074

T is transmission factor. R{nt is residual for internal agreement of symmetry
equivalent reflections. Rw = (Luj(F0 - Fc)2/ZwF02f-5 R = IXiF0\ - IFCII/ZIIF0II

four-circle diffractometer with filtered Mo-
Koc radiation (A = 0.7107 A). The intensity
data at 0 and 1.16 GPa were collected with
a Huber four-circle diffractometer with
graphite monochromated Mo-Koc radiation
(X = 0.7093 A), but the cell parameters
measured with the Picker diffractometer
were used for the structure refinements for
the purpose of consistency.

Unit-cell parameters were measured
with the procedure of King and Finger
(1979), whereby several reflections are
measured in eight equivalent orientations.
The range of 20 for all reflections was 1 8-
31° in order to avoid systematic errors that
result from comparing angular data from
different ranges (Swanson et al., 1985).

Intensities were measured for all acces-
sible reflections in a hemisphere of recipro-
cal space with sin0/A < 0.7. Corrections
were made for Lorentz and polarization
effects, crystal absorption by the diamond
and beryllium components of the pressure

cell. Digitized step data were integrated by
the method of Lehmann and Larsen ( 1 974).
Refinement conditions, refined isotropic
extinction coefficients, refined structural
parameters, and isotropic temperature fac-
tors are given in Tables 17 and 18.

As observed by Hazen et al. ( 1 99 1 ), the
c axis is the most compressible, with pc =
2.32(4) x 1 0-3 GPa- l . The a and b axes have
compressibilities of 1.72(14) x 10-3 and
1.71(3) x 10"3 GPa-1, respectively, — al-
most completely identical with the results
of Hazen et al. (1991), who reported the
value of 1.73(3)xl0-3 for the a and b axes
and 2.39(3) x 10"3 for the c axis. The
stiffness of a and b relative to c, according
to Hazen et al. (1991), results from the
pseudo-layering of the structure, with Mg-
octahedral layers parallel to (001 ), and cross
linking by Si207 tetrahedral pairs.

The pressure-volume data were fitted
with program VOLFIT to a Birch-
Mumaghan equation of state with 4 as the

GEOPHYSICAL LABORATORY

117

Table 18. Positional and equivalent isotropic thermal parameters.

Pressure, GPa

Parameter

1.16

1.81

2.88

4.84

Extn.

Coef(10-4)

0.37(5)

0.38(5)

0.26(5)

0.31(6)

0.26(6)

Ml, x,y,z

0.62(10)

0.66(11)

0.55(8)

0.45(7)

0.54(18)

M2,*

1/4

0.9696(5)

0.9694(6)

0.9704(4)

0.9697(4)

0.9687(9)

0.52(8)

0.43(6)

0.49(6)

0.47(13)

M3,*

1/4

0.1265(3)

0.1264(3)

0.1267(2)

0.1263(5)

1/4

0.91(6)

0.91(7)

0.70(5)

0.58(10)

Si, x

0.1203(2)

0.1200(2)

0.1199(1)

0.1202(3)

0.6166(2)

0.6165(3)

0.6167(2)

0.6168(2)

0.6172(4)

0.55(5)

0.54(5)

0.41(4)

0.37(4)

0.36(7)

Ol.z

1/4

0.2173(10) 0.2181(10)

0.2187(8)

0.2204(7)

0.2210(14)

0.9(2)

0.8(2)

0.5(1)

0.6(1)

0.2(2)

02,x

1/4

0.7158(9)

0.7167(9)

0.7159(8)

0.7158(7)

0.7199(14)

0.5(1)

0.5(2)

0.5(1)

0.3(2)

03,*

0.9889(5)

0.9895(5)

0.9900(4)

0.9899(4)

0.9900(8)

0.2564(8)

0.2557(8)

0.2554(6)

0.2553(5)

0.2538(15)

0.9(1)

0.6(1)

0.54(9)

0.40(8)

0.8(2)

04,*

0.2605(10) 0.2597(12)

0.2629(10)

0.2608(9)

0.258(2)

0.1226(3)

0.1225(4)

0.1218(4)

0.1227(2)

0.1211(7)

0.9923(4)

0.9933(5)

0.9933(4)

0.9925(3)

0.9936(7)

0.66(7)

0.71(8)

0.63(7)

0.58(6)

0.8(1)

assumed pressure derivative (K9). The bulk
modulus is 162(6) GPa if Vo is not con-
strained, comparable to the value 160(3)
GPa reported by Hazen et al. (1991).

The structure refinement of Fe-free
wadsleyite by Finger et al. (in preparation)
was used as the initial model in the present
work except that isotropic temperature fac-
tors were used in place of anisotropic ones
in their study. The space group, Imma, was
confirmed by the structure refinement. The
scattering factor curves for Mg, Si, and O

are those of neutral atoms in International
Tables for X-ray Crystallography (1974).

The RFINE6 program (Finger and
Prince, 1975) was used for the structure
refinements. The calculation converged
quickly to the R values listed in Table 17.
The extinction coefficients, atomic coordi-
nates, and isotropic temperature factors are
listed in Table 18.

Polyhedral volumes were calculated
using the program VOLCAL, written by
Finger (in Hazen and Finger, 1982); qua-

118

CARNEGIE INSTITUTION

Table 19. Polyhedral parameters for iron-free wadsleyite

Pressure, GPa
Parameter

0 1.16 1.81 2.88

4.84

Ml, Volume

QE*

M2, Volume

M3, Volume

Si, Volume

01, VS

02, VS

03, VS

04, VS

11.648(38)
1.0053(10)
15.83
2.17

11.880(38)
1.0056(46)
19.58
2.09

12.026(21)
1.0070(18)
23.11
2.06

2.276(10)
1.0037(40)
14.63
3.84

1.98
2.01
1.95
2.04

11.519(41)
1.0045(11)
13.62
2.21

11.779(42)
1.0061(51)
21.76
2.11

11.887(22)
1.0065(20)
21.74
2.10

2.294(11)
1.0034(43)
13.68
3.79

1.99
2.02
1.96
2.04

11.514(35)
1.0049(10)
15.52
2.21

11.959(35)
1.0057(39)
19.63
2.07

11.807(18)
1.0061(15)
20.10
2.12

2.254(9)
1.0033(35)
12.58
3.90

2.02
2.03
1.98
2.08

11.439(28)
1.0046(8)
14.43
2.24

11.768(30)
1.0056(34)
19.79
2.12

11.785(16)
1.0058(13)
18.79
2.13

2.254(8)
1.0040(30)
15.90
3.89

2.02
2.05
1.99
2.09

10.966(73)
1.0044(21)
13.35
2.39

11.554(70)
1.0065(79)
23.01
2.18

11.591(33)
1.0057(31)
18.72
2.19

2.298(20)
1.0026(71)
10.50
3.79

2.04
2.07
2.02
2.09

*QE and AV stand for the quadratic elongation and angle variance parameters, respec-
tively, of Robinson et al. (1969). VS stands for the valence sum of Brown (1981).

dratic elongations and angular variances
were calculated according to the defini-
tions of Robinson et al. (1971); the bond
valence sums are after Brown (1981). The
results are listed in Table 19.

The Si04 tetrahedral volume does not
decrease significantly with pressure, typi-
cal of all high-pressure studies below 10
GPa. Of the three Mg-0 octahedra, M3 is
the largest and most distorted at room pres-
sure (see Table 19); as pressure increases, it
becomes smaller and perhaps more regular.
Changes of the Ml and M2 octahedra are
less well defined (Fig. 64). The Ml and M2
octahedra become smaller with increasing

pressure, but there is little change in the
distortion.

Least-square fit of volume-pressure data
of the three polyhedra gives average
compressibilities of 8.7(11), 4.3(11), and
6.9(6) x 10'3 GPa"1, respectively, forM7,
M2, and M3 sites, corresponding to bulk
moduli of 116(15), 234(59), and 145(12)
GPa.

Zhang et al. (1991) demonstrated that
the unit-cell volume of an isostructural
series is approximately linearly related to
the bond length of a specific cation-anion
bond as long as other bond lengths do not
change. The compression of wadsleyite fits

GEOPHYSICAL LABORATORY

119

11.80

V= 11.677(30) -0.1 01 (13)P

K = 116(15) GPa
*

i i i i — i i i — i i i i i i i i i i i i — i — i — i i i

V=1 1.92(3) -0.051(13)P
I K = 234(59) GPa

11.60

11.50

£< 12.00

i i i — i i i i i i i

i i l i — i i i i — i— i-

V= 11.998(16) -0.083(7)P
K = 145(12) GPa

2.0 3.0 4.0

Pressure, GPa

5.0

Fig. 64. Variations of the volumes of the three M-
O octahedra with pressure.

into this specification because the Si-0
bond length is virtually constant. Figure
65 shows the variation of the unit-cell vol-
ume with the average of the three M-0
bond lengths; the linear relationship is a
clear indication that the contraction of the
unit cell is largely due to the shortening of
the M-O bonds. This linear relationship is
also true when Fe2+ substitutes for Mg2+ in
the wadsleyite structure and thus increases
the average M-0 bond length. In Fig. 66,
using the data of Finger et al. (in prepara-
tion), the unit-cell volume of wadsleyite
with different Fe2+ contents is plotted vs.
the average M-0 bond length. Again it

520

2.045 2.055 2.065 2.075 2.085

Average M-0 bond length, A

Fig 65. Linear relationship between the unit-cell
volume and the average of the three M-O bond
lengths.

co
©

■*±

■

546

Fe25

• S

544

•

542

Fe08v/

' Fe16

■

540

• yr

•

rqci

x^»

Te00

2.078 2.081 2.084 2.087 2.090 2.093

Average M-0 bond length, A

Fig 66. Unit-cell volume as a function of the
average M-0 bond length. The increases of the
average M-O bond length and the unit-cell volume
are due to the substitution of Fe2+ for Mg2+.

suggests that the increase or decrease of the
unit-cell volume is due to the increase or
decrease of the M-0 bond lengths.
We reach several conclusions
(1). Measurements of unit-cell param-
eters at 10 pressures give the axial
compressibilities of wadsleyite, which are
1.72(14), 1.71(3) x 10-3, and 2.32(4) x 10-
3 GPa-1, for a, by and c, respectively. As-
suming a Birch- Murnaghan equation of
state with K' = 4, the bulk modulus is
162(6) GPa.

120

CARNEGIE INSTITUTION

(2). The atomic coordinates, bond
lengths, bond angles, and polyhedral vol-
umes of the cation sites are documented.
The bulk moduli of the three octahedra are
116(15),234(59),andl45(12)GPaforM7,
M2, and Mi, respectively. M3 is the octa-
hedron for which the volume vs. pressure
data points are the least scattered and is also
the one which seems to become more regu-
lar with increasing pressure. The other two
octahedra, especially M2 , have bulk moduli
with large errors, and no significant varia-
tion in distortion is observed.

(3). The unit-cell volume is linearly
related to the average M-0 bond length, an
indication that the decrease of the unit-cell
volume is due to the shortening of the M-0
bonds.

References

Brown, I. D., The bond-valence method: an em-
pirical approach to chemical structure and
bonding, in Structure and Bonding in Crys-
tals, O'Keefe and Navrotsky, eds., 2, 1-30,
Academic Press, Boston, 1981.

Downs, J. W., Possible sites for protonation in (i-
Mg2Si04 from an experimentally derived elec-
trostatic potential, Amer. Mineral, 74, 1 124-
1129,1989.

Finger, L. W., andE. Prince, A system of Fortran
IV computer programs for crystal structure
computations, US National Bureau of Stan-
dard Technical Note 854, Washington DC,
1975.

Hazen, R. M., and L. W. Finger, Calcium fluoride
as an internal pressure standard in high pres-
sure/high-temperature crystallography,/. Ap-
plied Cry stallogr., 14, 234-236, 1981.

Hazen, R. M., and Finger, L. W., Comparative
Crystal Chemistry, Wiley, New York, 1982.

Hazen, R. M., J. Zhang, and J. Ko, Effects of Fe/
Mg on the compressibility of synthetic
wadsleyite: 0-(Mg \.x^x)2^0a (jc<0.25),
Phys. Chem. Minerals, 77, 416-419, 1991.

International Tables of X-ray Crystallography,
Vol. IV, Kynock Press, Birmingham, 1974.

King, H. E., and L. W. Finger, Diffracted beam
crystal centering and its application to high-
pressure crystallography, /. Applied
Crystallogr., 12, 374-378, 1979.

Lehmann, M. S., and M. K. Larsen, A method for
location of the peaks in step-scan-measured
Bragg reflections, Acta Cry stallogr., A30, 580-
584, 1974.

Robinson, K., G. V. Gibbs, and P. H. Ribbe,
Quadratic elongation: A quantitative measure
of distortion in coordination polyhedra, Sci-
ence, 772,567-570, 1971.

Smith, J. R., p-Mg2SiC>4: a potential host for
water in the mantle? Amer. Mineral., 72, 1051-
1055, 1987.

Swanson, D. K., D. J. Weidner, C. T. Prewitt, and
J. J. Kandelin, Single-crystal compression of
Y-Mg2Si04, EOS, 66, 370, 1985.

Zhang, J. M, D. N. Ye, and C. T. Prewitt, Rela-
tionship between the unit-cell volumes and
cation radii of isostructural compounds and
the additivity of the molecular volumes of
carbonates, Amer. Mineral., 76, 100-105, 1991.

Phase Transitions in
Framework Minerals

David Palmer

It is now clear that certain phase transi-
tions previously dismissed as being "subtle"
phenomena, may in fact produce dramatic
anomalies in the physical properties of
minerals and significantly modify their ther-
modynamic behavior. These phase transi-
tions, which typically involve displacive
distortions of the crystal lattice or cation
ordering effects, are common in most rock-
forming minerals that exist in the Earth's
crust. They are also expected to occur in
mantle phases.

Unlike heterogeneous reactions be-
tween crystallographically unrelated phases
(which are more usually studied by earth
scientists), the influence of a displacive or
order/disorder phase transition is not con-

GEOPHYSICAL LABORATORY

121

fined to the phase boundaries of an equilib-
rium system, but is significant at pressures
and temperatures far below the transition
point itself. As a consequence, stability
relations between mineral assemblages may
be perturbed throughout P-T space.

Quantitative Analysis of
Mineral Behavior

Most of the breakthroughs in the study
of phase transitions have come from solid-
state physics, the original motivation hav-
ing been to relate anomalies in physical and
thermodynamic properties to changes at a
crystal structural level. From a substantial
body of work on simple crystal structures,
came a number of "mean field" theories, all
relating to the macroscopic behavior of
crystals. These concern those phase transi-
tions which involve the lowering of crystal
symmetry, such that a relationship between
the symmetries of "high" and "low" phases
is maintained (i.e., a supergroup-subgroup
relation). Most displacive and order/disor-
der phase transitions fit into this category.

Macroscopic theories based on ideas
initially propounded by Landau (Landau
and Lifshitz, 1980) have been used exten-
sively in rationalizing the temperature de-
pendence of crystal behavior. The funda-
mental starting point for these theories is
the concept of a macroscopic order param-
eter, 2, which measures the progress of the
phase transition, such that Q = 0 in the high-
symmetry phase, and 0<Q<1 in the low-
symmetry phase. This might relate, for
example, to the ordering of magnetic spins,
positions of certain sets of cations such as

Al and Si, or to a prevailing lattice distor-
tion. Landau's original postulate was that
the energy lowering due to the high - low
transition, the excess free energy of the
low-symmetry phase, could be represented
as a power series in Q. From this potential,
it is possible to derive the temperature
dependence of the order parameter, in terms
of a critical exponent, p. It also becomes
possible to relate the excess thermody-
namic properties — heat capacity, entropy,
enthalpy etc. — to the order parameter and
hence to the progress of the phase transi-
tion. Further developments of Landau
Theory allow the behavior of excess physi-
cal properties to be related to the order
parameter and the thermodynamic proper-
ties.

The Landau approach, by relating all
excess physical, thermodynamic and struc-
tural properties to the macroscopic order
parameter considerably simplifies the de-
scription of phase transition behavior and
provides for a detailed understanding of
more complex, mineralogical systems.
Many minerals undergo more than one
phase transition, which may lead to seem-
ingly very complicated behavior: phase
transitions in the same material are rarely
independent of each other. In Landau
Theory, these effects are predicted in terms
of coupling between the various order
parameters, according to strict symmetry
rules.

The aim of this particular study is to
increase our understanding of phase-tran-
sition-related mineral behavior, focusing
on two less well studied areas, (1) high-
temperature thermodynamic properties, and
(2) high-pressure structural behavior.

122

CARNEGIE INSTITUTION

Framework minerals provide the "model
systems" for this work: these are the most
abundant materials at the Earth's surface;
their interconnected topologies ensure long
correlation lengths throughout the crystal
structure, and account for the many
displacive and cation-ordering phase tran-
sitions which occur with increasing tem-
perature and pressure. In this context there-
fore, "mean field" theories such as Landau
Theory are particularly relevant.

Order I Disorder Relations in Anorthite

Feldspars dominate the mineralogy of
the Earth's surface. They also show some
of the most complex subsolidus behavior
of all minerals, which may be attributed to
the interplay between the effects of succes-
sive displacive and cation-ordering phase
transitions. Much of this behavior has now
been quantified and rationalized within the
scope of Landau Theory (Carpenter, 1988).

The ordering of Al and Si at low tem-
peratures is common in many minerals, but
it has always been assumed that this pro-
cess is largely configurational, and does
not change the heat capacity. However,
many studies of feldspars have shown that
Al/Si ordering induces a lattice strain in the
crystal, thereby coupling to any prevalent
displacive distortion. It would be logical to
suppose that such ordering would
renormalize phonon frequencies, thereby
altering the heat capacity. If this does turn
out to be the case, then current attitudes to
mineral energetics will have to be reevalu-
ated.

One is not predicting substantial ACP
effects, and so testing this hypothesis re-
quires sensitive experimental procedures.
It would be desirable to compare the ener-
getics of two (or more) samples from the
same specimen, which differ only in the
extent of their Al/Si ordering (measured by
the order parameter Qod)- Changes in the
Al/Si ordering require solid-state diffusion,
which is an extremely sluggish process
below 1500 K or so. In order to be able to
prepare samples with different Qod in the
laboratory, one needs a material with a very
high equilibrium order/disorder phase tran-
sition temperature (Tc). Samples can then
be studied at lower temperatures without
the risk of continued ordering.

Anorthite, CaAl2Si20s, an end-mem-
ber plagioclase feldspar, is a good candi-
date for this study. Not only has the mineral
been extensively studied, its transition be-
havior is extremely well characterized (Car-
penter 1988, 1991). On heating, there is a
displacive phase transition from P\ to 7T at
510 K. Continued heating induces pro-
gressive disorder of Al and Si over the
tetrahedral sites, and the symmetry ap-
proaches CT, although anorthite melts (at
-1800 K) before reaching the hypothetical
/T - C\ phase transition temperature. The
most ordered samples available have Qod
~ 0.92, and this can be reduced, as required,
by annealing at high temperatures.

For this study on the effect of ordering
on the molar heat capacity, a natural anor-
thite was used (Qod = 0.92, as measured by
x-ray diffraction). A portion of the material
was annealed at 1723 K for 21 days to
induce some Al/Si disorder, thereby reduc-
ing Qod to 0.82. These samples were

GEOPHYSICAL LABORATORY

123

provided by Michael Carpenter (Univer-
sity of Cambridge). Differential scanning
calorimetry (DSC) runs performed in Cam-
bridge revealed no differences in the heat
capacities of these samples at low tempera-
tures (T< 900 K). However, this technique
cannot be used at higher temperatures and
so for a complete investigation it was de-
cided to use transposed-temperature drop
calorimetry at Princeton University, in col-
laboration with Alexandra Navrotsky. This
technique permits determination of heat
capacity, albeit indirectly, by measuring
the relative enthalpies at different tempera-
tures. Precise determination of Cp necessi-
tates many measurements at closely spaced
temperatures, but as a preliminary test to
see whether Al/Si ordering does induce a
ACp effect, it is sufficient to measure en-
thalpies at a few temperatures, to see if AH
between the samples varies as a function of
temperature.

Samples of anorthite were sealed in Pt
foil and dropped into a receptacle within a
"Setaram" calorimeter, set to the desired
temperature. The enthalpy change of the
sample from room temperature to the calo-
rimeter temperature is then proportional to
the energy required to restore the tempera-
ture of the assembly. An alumina-filled
capsule was used as calibration standard.
Enthalpy measurements were repeated three
to seven times per sample at each tempera-
ture, to check reproducibility. The results
are displayed in Table 20.

It must be stressed that these are rela-
tive enthalpies, that is, //r-//294K- The ac-
tual enthalpy of ordering between the
samples is ~ 8 kJmoH (Carpenter, pers.
comm.). The increase in AH between the

samples with increasing temperature is sig-
nificant, and may well indicate a ACp ef-
fect. Only at the highest temperature is
there the possibility of some slight disor-
dering within the calorimeter, though this
seems unlikely because of the short mea-
surement time (15 minutes, typically).
Because the samples used were natural,
albeit very pure, one cannot entirely dis-
count other effects due to trace impurities.
A second set of experiments, using syn-
thetic anorthites is now underway in order
to clarify this point.

These measurements have shown that
it is possible to measure enthalpy differ-
ences between minerals on the order of a
few joules for a 30-mg sample (3 kJmol-1
for anorthite) at temperatures far beyond
the reach of conventional calorimetric tech-
niques. Although this method is not as
precise as low-temperature DSC, further
enhancements, such as more sensitive de-
tector systems, are being developed to al-
low increased resolution at higher tem-
peratures.

Table 20. Relative enthalpy of anorthite from
973 K to 1673 K.*

T[K]

HtH294K

Qod=0M

[kJmol1]

Qoo=om

973

1273
1473
1673

194(4)
288(3)
349(3)
429(6)

195(3)
290(4)
369(3)
444(8)

* Errors are twice the standard deviation of
the mean.

124

CARNEGIE INSTITUTION

High-Pressure Studies of Phase
Transitions in Minerals

A New, High Pressure Phase Transition
in Leucite

It is possible to extend Landau Theory
to the description of high-pressure phase
transition behavior, using suitably chosen
"model systems". The aim of this research
is to concentrate on rock-forming minerals
that exist within the Earth's crust, where
moderate temperatures and pressures pre-
vail. The study of the pressure dependence
of such phases provides a logical step from
previous, high- temperature studies of phase
transition behavior.

Feldspathoid minerals were selected for
study. They are relatively abundant within
the crust and, like the feldspars, have alu-
minosilicate frameworks with cations in
interstitial sites. Phase transitions involv-
ing both displacive distortions and cation
ordering are extremely common.
Feldspathoids may be distinguished from
feldspars by the presence of structural chan-
nels, which provide an important reposito-
ries for cations, water, organic molecules
etc., and are pathways for ionic conduction
(Palmer and Salje, 1990; Alpena ai, 1977).
In this sense, feldspathoids may be com-
pared to zeolite minerals. The presence of
structural channels makes feldspathoids
low-density minerals; thus feldspathoids
might be expected to be very unstable at
high pressures. However, high-tempera-
ture studies of feldspathoids reveal a re-
markable structural adaptability, and this
may prolong the stability (or metastability)
of such structures at much higher pressures
than previously imagined.

Leucite, KAIS12O6, is a feldspathoid
mineral associated with SiC>2-poor, K-rich
alkaline volcanics, which has been well-
studied as a function of temperature. On
cooling, leucite undergoes two phase tran-
sitions, from a cubic phase Ia3d to an
intermediate tetragonal phase I4\lacd at Tc
= 938 K, described by order parameter Qi
(Eg symmetry); then to a low-7 tetragonal
phase 14 1 la at 7c = 918 K, described by
order parameter Qu with Tig symmetry
(Palmer er al, 1989). There is a continuous
volume decrease in the low-rphase, asso-
ciated with collapse of the <1 1 1> structural
channels; this is described by Qu . The first
order parameter, Qjy describes a ferroelastic
distortion of the unit cell (no volume
change). Such a distortion — an acoustic
shear mode — should show little or no
pressure dependence, compared to the
highly pressure dependent volume distor-
tion. Increasing pressure, therefore, is ex-
pected to modify the transition behavior by
enhancing Qu relative to Q/.

The pressure dependence of leucite has
been followed up to 60 kbar, using Raman
spectroscopy. A single crystal was placed
within a Mao-Bell diamond-anvil cell, us-
ing an organic liquid, "FC75" (C8F16O) as
pressure medium. Pressure determination
was by the ruby fluorescence method, with
ruby spectra measured using profile refine-
ment.

Measurement of the Raman spectra was
possible using a 0.6W He-Ne laser, with a

GEOPHYSICAL LABORATORY

125

580

570

I 560

| 550

c
®

| 540

530

520

■ increasing pressure
□ decreasing pressure

10 20 30 40
Pressure, kbar

Fig. 67. Pressure dependence of the Eg Raman
mode forleucite. The large increase in frequency
at Pc, and a hysteresis on increasing and decreas-
ing pressure, indicate the existence of a first-order
phase transition.

"Triplemate" detector and Princeton In-
struments control system. The leucite
Raman spectrum contains two intense
peaks, and a number of much weaker modes.
High-temperature Raman spectroscopy
(Palmer et aL, 1990) showed that the in-
tense modes relate to the two order param-
eters, and may be assigned to T\g and Eg
symmetries.

For this study, the pressure dependence
of the Eg mode was followed, revealing the
presence of a previously unknown phase
transition at P = 23 kb on increasing pres-
sure (Fig. 67). The similarity of the Raman
spectra on either side of the phase transi-
tion suggests that both "high" and "low"
phases are related. The fact that the phase
transition is totally reversible, together with
its rapid kinetics, suggest that the mecha-
nism is displacive. The existence of a
hysteresis implies that the phase transition
is first order in character. The existence of

a supergroup/subgroup relation to this phase
transition limits the choice of possible high-
pressure phases. We hope to carry out x-ray
work to refine the structure of the high
pressure phase, and to fully characterize
the phase transition behavior.

References

Alpen, U.U., H. Schulz, G. H. Tatat, and H.
Boehm, One-dimensional cooperarive Li-dif-
fusion in p-eucryptite. Solid State Commun.,
25,911-914, 1977.

Carpenter, M. A., Thermochemistry of alumi-
nium/silicon ordering in feldspar minerals, in
Physical Properties and Thermodynamic
Behaviour of Minerals, E.K.H. Salje, ed.,
D.Reidel, Dordrecht, pp. 265-323, 1988.

Carpenter, M. A., Thermodynamics of phase tran-
sitions in minerals: A macroscopic approach.,
in Stability of Minerals, G.D. Price, ed., Allen
and Unwin, Boston (in press), 1991.

Landau, L.D., and E. M. Lifshitz, Statistical
Physics, Edition 3, part 1, Pergamon Press,
Oxford, 1980.

Palmer, D.C., and E. K. H. Salje, Phase transitions
in leucite: dielectric properties and transition
mechanism. Phys. Chem. Minerals, 17, 444-
452, 1990.

Palmer, D.C., E. K. H. Salje, and W. W. Schmahl,
Phase Transitions in leucite: X-ray diffraction
studies. Phys. Chem. Minerals, 16, 714-719,
1989.

Palmer, D.C., U. Bismayer, and E. K. H. Salje,
Phase transitions in leucite: Order parameter
behaviour and the Landau Potential deduced
from Raman spectroscopy and birefringence
studies. Phys. Chem. Minerals, 17, 259-265,
1990.

Salje, E., B. Kuscholke, B. Wruck, and H. Kroll,
Thermodynamics of sodium feldspar II: ex-
perimental results and numerical calculations.
Phys. Chem. Minerals, 12, 99-1071, 1985.

126

CARNEGIE INSTITUTION

First-principles Studies of Elasticity

and Post-Stishovite Phase Transitions

in S1O2*

Ronald E. Cohen

Stishovite is a candidate mineral for the
Earth's transition zone and lower mantle,
and is also the prototypical octahedrally
coordinated silicate. It is an open ques-
tion, however, whether stishovite remains
the stable form of Si02 throughout the
lowermantle, or if a new structure for Si02
forms at the high pressure conditions of the
lower mantle. This is an important ques-
tion, because the presence or absence of
stishovite depends on the chemical compo-
sition of the lower mantle, particularly the
Fe-Mg ratio (e.g. Fei and Hemley, 1991).

Two recent studies suggestpossible phase
transitions under mantle conditions. Park
et al. (1988) predicted a phase transition
from stishovite (rutile structure) to the
pyrite structure (Pa3) at 60 GPa using self-
consistent Linearized Augmented Plane
Wave (LAPW) calculations. Tsuchidaand
Yagi(1989) looked for this transition using
in situ x-ray diffraction in the diamond
anvil cell, and instead reported a phase
transition in Si02 from stishovite to the
CaCl2 structure between 80-100 GPa. A
transition from stishovite to the CaCb struc-
ture in the lower mantle would be very
important in geophysical modeling of the
Earth. This phase transition involves an
elastic instability where c\\-c\2 becomes

*The computations were performed on the Cray 2
at the National Center for Supercomputing Appli-
cations under the auspices of the National Science
Foundation.

unstable (Cohen, 1987; Hemley, 1987).
Such a transition should be evident in seis-
mological data (derived acoustic veloci-
ties) if stishovite is present in any quantity
in the deep Earth. No such features are
observed in the lowermantle except for the
anomalous D" zone at the base of the
mantle. One must conclude either that
little stishovite is present in the deep man-
tle, or that the phase transition in Si02
occurs in D" and seismic anomalies in D"
reflect at least partially that transition.

These questions are addressed here us-
ing the Linearized Augmented Plane Wave
method (Wei and Krakauer, 1985). The
present calculations represent one of the
most extensive studies of a single material
using this technique — over 200 self-con-
sistent calculations were performed to de-
termine the phase relations, elasticity, and
vibrational properties of Si02 in the
stishovite, CaCl2, and Pa3 structures
(Cohen, 1 99 1 a,b). These calculations make
no assumptions about ionicity, bonding, or
form of the electron distribution. The only
inputs are the nuclear charges and posi-
tions, and the output is a total energy for
that nuclear configuration. The quantum
mechanical problem is solved within the
local density approximation (LDA) (Hedin
and Lundqvist, 1 97 1 ) of density functional
theory (Hohenberg and Kohn, 1964). The
electrostatic and kinetic energy contribu-
tions to the energy and potential are evalu-
ated numerically, and can be converged to
any necessary accuracy. In the LDA, the
quantum-chemical contributions are mod-
eled by assuming that the exchange and
correlation contributions to the potential
and energy can be obtained from the ex-

GEOPHYSICAL LABORATORY

127

O O*

stishovite

I i ■ i 'i

v(AJ)

Fig. 68. Equation-of- state (energy versus volume
at 0 K) of stishovite, CaCl2, and Pa3. The transi-
tion from stishovite to Pa3 is at about 160 GPa.
CaCh becomes stable relative to stishovite at 45
GPa. The initial energy difference between CaCl2
and stishovite is very small relative to the energy
changes with volume.

50 75

P (GPa)

Fig. 69. The elastic constant c\ \-c\2 as a function
of pressure. The solid line is a spline fit to the
calculated points (circles). The plus is the Brillouin
scattering data of Weidner et al. (1982). The
dashed curve is for the high pressure CaCl2 phase.
The phase transition is predicted to occur at 45
GPa at 0 K.

change and correlation functionals for the
uniform electron gas. During the last ten
years, this approximation has been demon-
strated to be quite accurate for metals,
semiconductors, and insulators, with some
exceptions primarily in magnetic crystals.

Calculations were performed for Si02 in
the stishovite (rutile, space group P4jJ
mnm), CaCl2 {Pnnm), and pyrite (Pa3)
structures as functions of volume. The
internal structural parameters and lattice
parameters were optimized at each vol-
ume. Fitting the total energies to an equa-
tion-of-state and to polynomials as func-
tions of strain and phonon amplitudes gives
the Alg and Big Raman modes as well as
four independent elastic constants for
stishovite as functions of pressure. The set
of calculations also gives the phase transi-
tion pressures from stishovite to the Pa3
and CaCl2 structures. All of the present
calculations are for static lattice energies,
classically equivalent to a temperature of 0
K.

Figure 68 shows the energy versus vol-
ume for the three phases. A phase transi-
tion from stishovite to Pa3 is predicted at
156 GPa. This is significantly higher in
pressure than the value of 60 GPa obtained
by Park et al. (1988) using LAPW. The
reason for the difference is not clear, but
extensive convergence tests of the present
results indicate stability of about 1 GPa in
the transition pressure with respect to the
convergence parameters.

Figure 69 shows the calculated elastic
constant c\\-c\2 as a function of pressure.
An elastic instability (c\\-c\2 vanishes) in
stishovite is predicted at 45 GPa, at which

128

CARNEGIE INSTITUTION

pressure stishovite would transform into
the CaCl2 structure at zero temperature,
Figure 68 shows that this phase transition
has only a small effect on the energetics of
Si02. It also has only a small effect, ini-
tially, on the crystal structure. The phase
transition is continuous, and the initial dis-
tortions from the rutile structure are very
small. However, the phase transition has
enormous effects on the elastic properties
of high pressure Si02, so it is crucial to
consider whether this transition occurs un-
der mantle conditions.

The phase transition is likely to be very
sensitive to temperature since the instability
is driven by a Raman mode. Figure 70
shows the Aig and Big Raman modes as
functions of pressure. Agreement with the
Raman data obtained by Hemley (1987) is
excellent. At the CaCl2 phase transition
the Big mode becomes an Ag mode and
begins to increase in frequency with pres-
sure, whereas as lower pressures the fre-
quency decreases with increasing pressure.
This is probably the most direct way of
detecting the phase transition, since the
distortions are very small at the transition
and would be difficult to detect with x-rays.
The phase transition is not a soft-mode
transition in the traditional sense, since the
Big frequency does not reach zero at the
transition. The Big mode does drive the
transition, since if the atoms were not al-
lowed to distort along the Big eigenvector
during a c\i-c\2 strain, c\\-c 12 would not
become unstable. It is the coupling be-
tween the strain and the phonon displace-
ment that makes the energy decrease as a
function of strain, and thus causes the phase
transition. The Big Raman mode contrib-

150

Fig. 70. Pressure dependence of the Aig and Big
Raman frequencies for stishovite, and an Ag
mode in CaCl2- The triangles and diamonds are
experimental data from Hemley (1987). The
dashed line is a prediction for the CaCl2 phase.
Note that the "soft mode" does not go to zero, or
even get very small at the phase transition.

utes a factor to c\\-c\2 that is proportional
to -1/co2 (Miller and Axe, 1967), so that as
the Big frequency decreases the elastic
constant c\\-c yi is destabilized. Tempera-
ture, however, would stabilize the higher
symmetry rutile structure. At 150 GPa, the
well depth (energy gain on the phase tran-
sition) expressed as temperature is only
1 275 K, so at mantle temperatures of 2000-
3000 K the transition may not occur until
much higher pressures than the 45 GPa
calculated for zero temperature. The exact
transition temperature can be calculated
using molecular dynamics or Monte Carlo
simulations using a potential model, or by
fitting a simplified Hamiltonian to the total
energy results and using statistical thermo-
dynamics to evaluate the phase diagram
(Rabe and Joannopoulos, 1987). Both ap-
proaches require information about the

GEOPHYSICAL LABORATORY

129

coupling between cells, as well as the total
energies for zone center distortions that do
not increase the unit cell size. A potential
model is now being developed for high-
pressure Si02 to investigate the thermal
properties of stishovite and the tempera-
ture dependence of the phase transition. It
is quite possible that temperature will sig-
nificantly increase the depth at which the
transition occurs, and thus the transition
may be partly responsible for the anoma-
lous seismic properties of the so-called D"
region.

References

Cohen, R. E., Calculation of elasticity and high
pressure instabilities in corundum and
stishovite with the potential induced breathing
model, Geophys. Res. Lett., 14, 37-40, 1987.

Cohen, R. E., Bonding and elasticity of stishovite
Si02 at high pressure: Linearized augmented
plane wave calculations, Amer. Mineral., 76,
733-742, 1991a.

Cohen, R. E., First-principles predictions of elas-
ticity and phase transitions in high pressure
Si02 and geophysical implications, in High
Pressure Research in Mineral Physics: Appli-
cation to Earth and Planetary Science (Pro-
ceedings of U.S. -Japan Conference on High
Pressure Geophysics, Ise, Japan, January,
1991), M.H. Manghnani and Y. Syono, eds.,
in press, 1991b.

Fei, Y. and R. J. Hemley, Stability of (Fe,Mg)Si03-
perovskite in the lower mantle, Geophys. Res.
Lett., in press, 1991.

Hedin, L. and B. I. Lundqvist, Explicit local
exchange-correlation potentials, /. Phys., C4,
2064-2083, 1971.

Hemley, R. J., Pressure dependence of Raman
spectra of Si02 polymorphs: a-quartz, coesite,
and stishovite. In High -Pressure Research in
Mineral Physics, M.H. Manghnani and Y.
Syono, eds., pp. 347-359. American Geo-
physical Union, Washington, D.C., 1987

Hohenberg, P., and W. Kohn, Inhomogeneous
electron gas, P/ry.s./?ev.,7.?<5£, 864-871, 1964.

Kohn, W. and L. J. Sham, Self-consistent equa-
tions including exchange and correlation ef-

fects, Phys. Rev., 140 A, 1133-1140, 1965.

Miller, P. B. and F. D. Axe, Internal strain and
Raman active vibrations in solids, Phys. Rev.,
163, 924-926, 1967.

Park, K. T., K. Terakura, and Y. Matsui, Theoreti-
cal evidence for a new ultra-high-pressure
phase of Si02, Nature, 336, 670-672, 1988.

Rabe, K. M., and J. D. Joannopoulos, Theory of
the structural phase transition of GeTe, Phys.
Rev. B, 36, 6631-6639, 1987.

Tsuchida, Y. and T. Yagi, A new, post-stishovite
high-pressure polymorph of silica, Nature,
340, 217-220, 1989.

Tsuneyuki, S . , M. Tsukada, H. Aoki, and Y. Matsui,
First-principles interatomic potential of silica
applied to molecular dynamics, Phys. Rev.
Lett., 61, 869-872, 1988.

Tsuneyuki, S., Y. Matsui, H. Aoki, andM. Tsukada,
New pressure-induced structural transforma-
tions in silica obtained by computer simula-
tion, Nature, 339, 209-21 1, 1989.

Wei, S. H., and H. Krakauer, Local density func-
tional calculation of the pressure induced phase
transition and metallization of BaSe and BaTe,
Phys. Rev. Lett., 55, 1200-1203, 1985.

Weidner, D. J., J. D. Bass, A. E. Ringwood, and W.
Sinclair, The single-crystal elastic moduli of
stishovite, /. Geophys. Res., 87, B4740-4746,
1982.

Molecular Dynamics Simulations of
Melting of MgO at High Pressures*

Zhaoxin Gong, Ronald E. Cohen, and
Larry L. Boyer**

We are developing a general-purpose
molecular dynamics (MD) program for
studying finite clusters of atoms using
many-body potentials and long-range
forces. This program will be used to study
high-pressure and high-temperature phase

* This work is supported by the Office of Naval
Researchgrant#N00014-91-J-1227toREC. Com-
putations were performed with the support of
ONR at the NRL Connection Machine Facility
and at the Pittsburgh Supercomputer Center under
the auspices of the National Science Foundation.
*Complex Systems Theory Branch, Naval Re-
search Laboratory, Washington, D.C. 20375

130

CARNEGIE INSTITUTION

transitions as well as ferroelectricity.
Though many simulations have been per-
formed on various systems of interest, rela-
tively few have been performed on ionic
systems using first-principles, non-empiri-
cal potentials. Such potentials are expected
to be more reliable outside the range of
experimental data than empirical poten-
tials, and furthermore can be constrained
for arbitrary ionic configurations, not
merely determined from bulk properties
with atomic positions close to their equilib-
rium sites. A cluster approach is important
because periodic boundary conditions have
been found to greatly inhibit phase transi-
tions such as melting, especially in systems
with long-range forces, because the peri-
odic boundary conditions force the "liq-
uid" state onto a periodic lattice. Also, in
the case of ferroelectricity, electric fields
and macroscopic polarization are by defi-
nition "non-periodic," so that these phe-
nomena can only be studied approximately
using periodic boundary conditions. Finite
clusters in free space were used in earlier
MD studies of melting of NaF using a
Gordon-Kim rigid ion potential (Boyer and
Pawley, 1988; Boyer and Edwardson,
1988).

In order to use a cluster method, many
atoms must be be included in order to
minimize the effects of surfaces; however,
with long-range forces the MD problem is
an N2 problem, where N is the number of
atoms in the cluster, so computational effi-
ciency is very important. We present re-
sults here using a massively parallel com-
puter, the Connection Machine CM-2, as
well as results calculated on an IBM RS/
6000 superworkstation and a Cray YMR

Here we report for the first time MD results
of melting of MgO at high pressures using
a many-body potential, the Potential In-
duced Breathing (PIB) model.

The PIB model is an ab initio model
where no parameters are fitted to the ex-
periment. This model has been very suc-
cessful in predicting the thermodynamic
and elastic properties of alkaline earth ox-
ides, including the Cauchy violations (c\\
is not equal to cyi at zero pressure) (Boyer
era/., 1985; Mehle/tf/., 1986; Cohen etaL,
1987; Isaak et al., 1990). Monte Carlo
(MC) and MD simulations for the PIB
model are computationally demanding and
complicated. So far only primitive MC
results of equation of state at zero pressure
using the PIB model have been reported for
a sample of MgO with 64 atoms (Cowley et
al„ 1990).

Isaak et al. (1990) found that the PIB
model predicts that the zero pressure iso-
thermal elastic modulus Cs(=c\i-C22) of
MgO becomes negative at a temperature
very close to the melting temperature of
MgO. The Cy instability occurs before the
bulk modulus instability discussed by Boyer
(1985) with increasing temperature if the
free energy, including the vibrational con-
tribution, is calculated as a function of
strain. To find out whether this instability
is related to the melting, we investigate
here whether the same PIB potential indi-
cates a melting point close to the
quasiharmonic elastic instability. Further-
more, are there elastic instabilities that cor-
respond to melting at high pressures?

An important question of geophysical
importance is the curvature of the melting
curve at high pressures. At low pressures

GEOPHYSICAL LABORATORY

131

the melt is generally (but not always) less
dense than the crystalline form of a mate-
rial, but the liquid is more compressible.
As pressure is increased, the difference in
densities of liquid and solid should de-
crease, and the slope of the melting curve,
dT/dP, should decrease with pressure. The
melting curve may become horizontal or
may reach a maximum and bend over.
MgO, being a close-packed solid, is an
interesting case because of the simple crys-
tal structure and the probably simple struc-
ture of the liquid. MgO is also an
endmember of magnesiowustite (Mg,Fe)0,
which is considered a likely mineral in the
Earth's lower mantle.

We used the same numerical PIB poten-
tials as in Isaak et al. (1990), but we have
fit them to a different form in order to
increase accuracy and computational effi-
ciency. One problem is that the Madelung
(electrostatic) potentials on the oxygen ions
on the surface of the cluster are lower in
magnitude than in the bulk. The quantita-
tive results presented here must be consid-
ered preliminary, because the surface atom
potentials often fell outside the range of the
points at which the PEB potential was fit.
We expect the changes to be small when a
better potential is used, because the form
we chose appears to extrapolate smoothly,
and the great majority of atoms had poten-
tials that fell within range of the fit.

We tested the MD code by comparing
the zero-pressure density of a free cluster
with that from quasiharmonic lattice dy-
namics (LD) calculations using the same
model (Isaak etai, 1990). The density was
calculated using the method of Boyer and
Pawley (1988). At T= 1300 K, the average

density for the innermost region of a cluster
with 216 atoms differs from the LD result
by about 7%. The difference decreases to
about 6 % when the cluster size is increased
to 512 atoms. We also found that the total
energy is conserved to better than six sig-
nificant figures over many thousands of
time steps, giving confidence in the MD
code.

We performed a series of runs to find
out the melting temperature at zero pres-
sure. The starting configuration of 216 (=
6x6x6) atoms in their perfect lattice
positions was initially given a small amount
of kinetic energy, equivalent to a tem-
perature T = 300 K. The simulation was
then allowed to proceed at constant energy
for 30 ps (10,000 time steps); then the total
energy of the cluster was increased by
scaling up the velocities to give an increase
in temperature of 400 K. This procedure
was repeated several times. The results are
drawn in Figures 71 and 72 as solid lines.

0.55

060 0.65

E + 275 (Hartree)

0.70

Fig. 71. Plot of equilibrium temperature T versus
the total energy E at P = 0 GPa. The solid line
corresponds to increasing energy while the dotted
line corresponds to decreasing energy.

132

CARNEGIE INSTITUTION

055

060

065

0.70

E + 275 (Hartree)

Fig. 72. Zero pressure plot of equilibrium density
versus the total energy, E, in a cubic region cen-
tered at the center of mass with the cube side equal
to the nearest neighbor distance of the perfect
lattice. The solid line corresponds to increasing
energy while the dotted line corresponds to de-
creasing energy.

At low temperatures, the temperature is
almost a linear function of the energy E.
Thus MgO is quite harmonic up to about
half the melting temperature. The tempera-
ture increases as the energy increases until
T= 3815 K, where increasing energy leads
to a temperature decrease to about 3045 K.
(This run is three times longer than the
other runs in this figure.) The temperature
of the cluster then increases again as the
energy is increased. The density of cluster,
on the other hand, decreases slowly as the
energy is increased until the temperature
reaches about T= 38 1 5 K. The density then
decreases by a much larger amount as the
energy is increased by another increment
and the temperature drops to 3045 K. This
simultaneous larger decrease in tempera-
ture and density indicates that the cluster
melted and kinetic energy is transferred
into potential energy (Boyer and Pawley,

1988; Boyer and Edwardson, 1988). The
dotted lines in these two figures correspond
to the process of cooling. With decreasing
energy, the cluster freezes to a crystal with
defects due to the high cooling rate, which
leads to hysteresis. The middle of the
hysteresis loop is about 3100 K, which
compares remarkably well with the experi-
mental melting temperature of 3098 ± 20 K
(Stull and Prophet, 1971) and the c\\ -c\ 2
elastic instability (Isaak et al., 1990) using
the same potential.

In a second series of calculations, the
cluster is confined in a box so that its
pressure can be varied. It is natural to
repeat procedures for a free cluster to find
kinks in the T versus E curve, while the
volume is kept fixed. At large volumes
(low pressures) this procedure is satisfac-
tory, but it becomes increasingly difficult
to locate the kinks as they becomes less and
less prominent as volume becomes smaller
and smaller. At a kink, either a large differ-
ence in temperature Tor a large difference
in pressure P is observed. When the cluster
melts, its temperature decreases while its
pressure increases. For example, when the
box size is set equal to 23.735 bohr, the
drop in temperature is about 260 K, from
about 5880 K to about 5620 K, and the
increment in pressure is about 4.5 GPa,
from about 7.7 GPa to about 12.2 GPa.
However, for a box of cube side of 23.23
bohr, we were not able to identify any kink
with the drop in temperature larger than 70
K and the increment in pressure larger than
1.7 GPa.

To determine the melting temperatures
at high pressures, we first find the tempera-
ture and energy at given pressure and given

GEOPHYSICAL LABORATORY

133

* 8-

0.55 0.60 0.65 0.70

E + 275 (Hartree)

0.75

Fig. 73. Plot of equilibrium temperature T versus
the total energy E at P = 10 GPa. The solid line
corresponds to increasing volume while the dotted
line corresponds to decreasing volume.

P (GPa)

Fig. 74. Plot of the melting temperature Tm versus
the pressure P. The bars indicate the maxima and
minima of the heating and cooling curves, respec-
tively, around the melting points.

volume, then we plot the temperature ver-
sus energy at a given pressure (Fig. 73).
The solid line represents the process of
increasing volume, while the dotted line
represents decreasing volume. The cube
side increment in this figure is 0.505 bohr,
and the starting box side is 20.2 bohr. We
identify the kinks in the plot as associated
with the melting.

In Figure 74 we plot the melting tem-
perature versus pressure. The melting tem-
perature, because of the van der Waals loop
typically present in our results, is taken to
be the average temperature of two middle
points in the heating and cooling curves.
Although only three points have been cal-
culated, it is obvious that there is signifi-
cant curvature in the melting curve. Melt-
ing temperatures are presented only to 20
GPa because of an instability that occurred
at high temperatures, due to the fitting
range of the Watson sphere potential. At
high temperatures the Watson potentials of

many atoms lie outside the fitting range,
which can be corrected by the development
of better inter-atomic potentials. Further
results will be on very high pressures and
larger samples.

Although the PIB model is more com-
plicated than the rigid ion model, many
calculations can utilize parallel processing.
For example, checking for atoms that cross
the box boundary can be done in parallel in
all three directions for all the atoms in the
cluster. The most time consuming part, the
force calculation, can also be done in paral-
lel. Many simulations were done on a CM-
2 from Thinking Machine Inc. at the Naval
Research Laboratory (NRL). On the CM-
2, the row column difference (RCD) method
by Boyer and Pawley (1988) is used when
calculating pair interactions. The block
sizes used are the largest possible, i.e, block
sizes of N x N, which has been shown to be
the most efficient way of calculating pair
interactions for a cluster of 512 atoms

134

CARNEGIE INSTITUTION

Table 21. Timing results (sec per time step) for MD simulations on the CM (one 8k
sequencer), the CRAY-YMP (one processor), and the IBM RS/6000-320 power worksta-
tion.

Rigid Ion Model*

PIB Model

CRAY

IBM

CPU time
Elapsed Time

216 atoms
512 atoms
1000 atoms

0.23
0.76

0.32
1.44

0.13
0.31
1.02

0.23
0.43
1.22

0.055
0.29
1.01

0.082
0.46
2.16

0.31

2.45
12.3

0.32
2.50
12.5

The timing results for rigid ion model are from Boyer and Edwardson (1988).

(Boyer and Edwardson, 1988). For a clus-
ter of 1000 atoms, Boyer and Edwardson
(1988) were limited to a block size of 256
x 256 by memory constraints. In the
present hardware at NRL, we encountered
no memory constraints for clusters of up to
1000 atoms. Timings are given in Table 21 .
The elapsed times are for an otherwise
empty machine except for the Cray tim-
ings, which were taken under a typical
load.

In summary, we have presented first
results on melting of a oxide using a non-
empirical many-body model. MD simula-
tions were performed on finite clusters
containing up to 1000 atoms on the mas-
sively parallel CM-2. Significant curva-
ture in the melting curve is predicted. This
method will be applicable to study a large
variety of phase transitions as functions of
temperature and pressure.

References

Boyer, L. L., Theory of melting based on lattice
instability, Phase Transitions, 5, 1-48, 1985.

Boyer, L. L., M. J. Mehl, J. L. Feldman, J. R.
Hardy, J. W. Flocken,, C. Y.andFong, Beyond
the rigid ion approximation with spherically

symmetric ions, Phys. Rev. Lett., 54, 1940-
1943, 1985.

Boyer, L. L., and P. J. Edwardson, Application of
massively parallel machines to molecular dy-
namics simulation of free clusters, Proceed-
ings of the 2nd symposium on the frontiers of
. massively parallel computation, Fairfax, Vir-
ginia, USA, October 10-12, 1988.

Boyer, L. L., andG. S. Pawley, Molecular dynam-
ics of clusters of particles interacting with
pairwise forces using a massively parallel com-
puter, /. Comput. Phys., 78, 405-423, 1988.

Cohen, R. E., L. L. Boyer, and M. J. Mehl, Lattice
dynamics of the potential-induced breathing
model: Phonon dispersion in the alkaline -earth
oxides, Phys. Rev. B, 35, 5749-5760, 1987.

Cowley, E. R., S. H. Liu, and G. K. Horton, Monte
Carlo calculations of the equations of state of
alkaline earth oxides, Ferroelectrics, 111, 33-
42, 1990.

Isaak, D. G., R. E. Cohen, and M. J. Mehl, Calcu-
lated elastic and thermal properties of MgO at
high pressures and temperatures, J. Geophys.
Res., 95, 7055-7067, 1990.

Mehl, M. J., R. J. Hemley, and L. L., Boyer,
Potential-induced breathing model for the elas-
tic moduli and high pressure behavior of the
cubic alkaline-earth oxides, Phys. Rev. B, 33,
8685-8696, 1986.

Stull, D. R. and H. Prophet (Eds.), JANAF Ther-
mochemical Tables, 2nd ed., Office of Stan-
dardReference Data, NIST, Washington, D.C.,
1971.

GEOPHYSICAL LABORATORY

135

Glass Diffraction Measurements with
Polychromatic Synchrotron Radiation

Charles Meade and Russell J. Hemley

The structure of liquids and glasses at
high pressures is an issue of great impor-
tance in the Earth and Material Sciences. At
present, however, little is known on this
subject because of the difficulty of obtain-
ing direct information about the structure
of noncrystalline materials at high pres-
sures. Spectroscopic studies (both optical
and x-ray) have provided constraints on the
vibrational properties and short-range or-
der in glasses under pressure, but they
typically provide only an indirect probe of
structure. Perhaps the most direct means of
determining glass and liquid structures at
very high pressures (P > 10 GPa) is to
obtain x-ray diffraction measurements from
these materials in the diamond cell. Here,
we investigate the use of high-energy poly-
chromatic synchrotron radiation for this
purpose.

To obtain precise measurements of glass
diffraction under these conditions, two prob-
lems must be addressed. First, one has to
know the intensity distribution of the x-ray
source to interpret the x-ray patterns. Sec-
ond, in the analysis of the measured dif-
fraction spectra, one must be able to sub-
tract the background diffracted intensity
produced by the diamond anvils (both Bragg
and Compton scattering) from the signal
due to the amorphous material.

In this report, we will address the first
question, constraining the x-ray source spec-
tra in glass diffraction experiments. To il-
lustrate a new analysis that we have devel-

oped for this problem, and to demonstrate
the advantages of using high energy syn-
chrotron radiation for these experiments,
we have measured the x-ray diffraction
from a small platelet of Si02 glass under
ambient pressures and temperatures at the
X-17C beamline of the National Synchro-
tron Light Source, Brookhaven National
Laboratory (B adding et a/., 1990; Mao et
al. 1990). Because these measurements
were made outside of the diamond cell on
a known material (Mozzi and Warren, 1 969;
Konnert and Karle, 1973), they provide an
initial test of our ability to constrain the
intensity distribution of the x-ray source in
our experiments.

Examples of the measured and known
diffraction patterns for Si02 glass are shown
in Figure 76. From this comparison, it is
immediately evident that that the shape of
the x-ray source distribution strongly influ-
ences the observed x-ray measurements.
To normalize these diffraction patterns,
one can measure the intensity of the x-ray
source; however, these measurements are
difficult given the the detector geometry at
X-17C. One could calculate the source
distribution (e.g. Krinsky et al., 1983),
though, this requires precise knowledge of
the upstream filters and the critical energy
of the superconducting wiggle r. Moreover,
calculated spectra cannot account for par-
tial absorption of the beam by upstream
slits.

Thus, we have adopted an alternative
approach where we take advantage of the
redundancy in our diffraction spectra ob-
tained at several different scattering angles.
Consider that the observed intensity spec-
tra at particular diffraction angle 20/ (cor-

136

CARNEGIE INSTITUTION

J tMJM TH"MJTint "" Jlttip I I I [I I H |Mtl|l

s.A-1

Fig. 76. (A) Diffraction pattern for Si02 glass at ambient pressures and temperatures measured from
monochromatic x-ray source by Konnert and Karle (1973). The horizontal axis is the wavenumber
S=47tsm0/X. (B) Observed intensities (corrected for sample and air absorption) from Si02 glass
measured at two different scattering angles (26) with a polychromatic x-ray source at beamline X-17C
of the National Synchrotron Light Source. The differences between these data and with the previously
measured values are due to the distribution of x-ray intensities in the polychromatic synchrotron source.

rected for air and sample x-ray absorption common scale we define the wavenumber
and polarization of the diffracted beam) s = AKsin6lX. We can then write equation
can be represented as (1 ) as

/$s(£) = [B(28i, E) + C(2Q, E)]f0(E) ( i ) /obsW = [B(s) + C(j)]/o(£,20)

(2)

where B and C are functions that respec-
tively describe the coherent (Bragg) and
incoherent (Compton) scattering from the
sample and Io is the source
spectrum. Whereas 5 and Care functions of
2 0 and the energy of the incident beam (E),
Io is a function of E only. In practice, we
make measurements at several diffraction
angles. Thus, to express all of the data on a

Here, B, C, and Iobs are functions of s only.
In this representation, the source function
depends on s and 20. With this transforma-
tion, the measurements at different diffrac-
tion angles can be collapsed down to the
single function Iobs(s). Even though/? and
C are not known a priori, one can use the
knowledge that they are functions of s only
to constrain the source function Iq. Specif i-

Fig. 77. (A) Composite diffraction pattern for Si02 from measurements at seven angles ranging from
2 0 = 6° to 45°. Below s = 1 , the data are extrapolated to zero intensity. (B) The same diffraction spectra
as shown in Figure 76B. When the spectra are normalized with the correct source function, they are
equivalent when plotted in terms of the wavenumber s. For clarity, the spectra are offset in the vertical
direction. (C) X-ray source function for X-17C that is consistent with these measurements. The
intensities drop off below 12 keV because of absorption in upstream beryllium and carbon filters.

GEOPHYSICAL LABORATORY

137

1""

1 1 ii

nrpiii|iiii|iin

1 1 ■ 1 1 1 1 1 ( 1

iTrrpTTT

Mill

•4— •

u„J

■ ml

Mill

jj,Liiii lui liin '

.ml....

I....I

111 1I111 1

7177

TTTTT

cally, we determine Io by requiring that
I0DS{s) must be the same for all diffraction
angles.

An example of a "composite" diffrac-
tion pattern obtained in this way and the
corresponding source function is shown in
Figure 77. When compared to the previ-
ously measured pattern, the advantages of
using this approach is apparent. Because
one can measure diffraction to high ener-
gies (-60 keV) at high angles, diffraction
spectra can be obtained to much higher
values of s than in conventional laboratory
measurements. In this study, we obtained
data to s = 26 A"1. Measurements to 40 A"
1 are possible. We expect that such mea-
surements over a wide range of s values
will allow strong constraints on glass struc-
ture at high pressures.

References

Badding, J.V., H. K. Mao, J. Z. Hu, R. J. Hemley,
C. Meade, and J. F. Shu, High pressure energy
dispersive x-ray diffraction at X-17C, EOS,
71, 1620, 1990.

Konnert, J. H., and J. Karle, The computation of
radial distribution functions for glassy materi-
als, Acta. Cryst., A29, 702-710, 1973.

Krinsky, S., M. L. Perlman, and R. E. Watson,
Characteristics of synchrotron radiation and
of its sources, in Handbook on Synchrotron
Radiation, E. E. Koch, ed., North Holland,
New York, pp. 65-171, 1983.

Mao, H. K., J. Z. Shu, J. F. Shu, and R. J. Hemley,
Ultrahigh-pressure experimentation above 300
GPa, Bull Ame.r Phys. Soc, 36, 529, 1990.

Mozzi, R. L., and B. E. Warren, The structure of
vitreous silica, /. Appl. Cryst., 2, 162-172,
1969.

10 20 30 40 50
Energy, keV

60 70

138

CARNEGIE INSTITUTION

X-ray Diffraction of Solid Nitrogen-
Helium Mixtures*

Willem L. Vos, Larry W. Finger,
Russell J. Hemley, Ho-Kwang Mao,

Jing Zhu Hu, Jin Fu Shu,

Richard LeSar* Andre de Kuijper, *

and Jan A. Schouten**

The study of simple molecular mixtures
in their solid phases at high pressures has
recently become of increasing interest. This
work is important for statistical and con-
densed matter physics, specifically to in-
vestigate the influence of size ratio, pack-
ing, and van der Waals-like interactions on
the stability of mixed structures. Such stud-
ies also provide a starting point for for
modeling the interiors of the outer planets
of our solar system.

One of the most studied mixtures in this
respect is N2-He; helium is still a fluid
under conditions where nitrogen has so-
lidified (see Fig. 78). At room temperature,
their freezing pressures differ by a factor of
5: i.e., 24 vs. 120 kbar). By a combination
of visual observations and quasi-isochoric
p-T scans, Vos and Schouten (1990) found
that around 10 mol% He is soluble in the
orientationally ordered rhombohedral e -
phase of nitrogen (Mills et ai, 1986); in
contrast, the solubility of helium in the
disordered (3 or 8 - phases (Cromer et ai,

*This collaboration was supported by a NATO
Collaborative Research Grant.
* T-ll, Los Alamos National Laboratory, Los
Alamos New Mexico 87545

van der Waals-Zeeman Laboratorium,
Universiteit van Amsterdam, 1018XE Amsterdam,
The Netherlands.

1981) is negligible. Due to this solubility,
the stability of the e-phase was found to
increase drastically with respect to the 8 -
phase (see Fig. 78). The shift of the 8 - e

100

J*
<xi

k—
CL

250 350 450 550

Temperature, K

Fig. 78. PTphase diagrams of N2 (solid lines), He
(short dashed line), and projected phase diagram
of N2-He (dashed lines). The solid lines are in
order of increasing pressure the P-fluid, 5-p and
e-8 lines, the short dashed line is the melting line
of helium and the dashed lines are the three - phase
lines e-8 -F of the mixture.

transition was subsequently confirmed by
vibrational Raman spectroscopy
(Scheerboom and Schouten, 1991). Here,
we report preliminary investigations of the
structure of the mixed phases (indicated by
asterisks) by x-ray diffraction and a com-
parison of the results with computer simu-
lations.

X-ray diffraction studies were performed
at beamline X-17C of the National Syn-
chrotron Light Source, at Brookhaven Na-
tional Laboratory, using a single-crystal
type diamond anvil cell (Mao and Bell,
1980). The samples were prepared from
high purity (99.999%) gases and loaded in

GEOPHYSICAL LABORATORY

139

a pressure vessel at a pressure of about 2
kbar. All experiments were performed at
room temperature, and pressures were ob-
tained from the linear ruby scale. Three
experiments were performed: one on pure
N2, to check the structure of the e-phase,
since this had been obtained from powder
diffraction (Mills et al., 1986, Olijnyk,
1990), and two on the 8* - phase at compo-
sitions of 5.0(2) and 10.0(2) mol% helium.
The diffraction experiment on e-N2 was
performed at 195 kbar since the transition
from 8 to e takes place at 1 65 kbar (Olijnyk,
1990). The sample showed clear colors
under crossed polarizers, which distin-
guishes it from the non-birefringent 8 -
phase. The sample consisted of at least one
large crystal and many small grains, since
strong reflections were observed at distinct
(X,oS) angles, while a powder-like pattern
with preferred orientation was observed
irrespective of the ix,(o) angles. This con-
figuration was often encountered and was
named "single-powder." The d-spacings of
the powder pattern agree with the ones
reported previously (Mills et aL, 1986;
Olijnyk, 1990) and yield a unit cell of
dimensions a= 7.63(4) A, c= 10.37(10) A,
c/a= 1.36 with 24 molecules and space

group R3c (Mills et ai, 1986). Ten reflec-
tions were observed from the single crys-
tal, that were also consistent with the R3c
structure. Special attention was paid to
overtones and to possible lower-order re-
flections, since some computer simulation
results yielded superstructures with unit
cells containing up to 64 molecules (Nose
and Klein, 1986, Belak era/., 1990). How-
ever, no longer spacings were found and
only a few weak overtones were observed,

also consistent with the aforementioned
structure.

With 5 mol % He, experiments were
performed at 126 and 144 kbar. The pres-
ence of the e* - phase was checked by the
observation of birefringence. The sample
consisted again of a "single-powder." The
spectrum of d-spacings consisted of lines
from both 8* and £* - phases. This was
confirmed by Raman scattering, which
showed that the vi mode was clearly split,
which is not the case in the pure phases. The
spacings of the 8* - phase are the same as
those of the 8 - phase at the same pressures
(Olijnyk, 1990), which means that there is
negligible solubility of helium in this phase.
The other d-spacings could all be fitted to a
hexagonal phase with a - 8.050(5) A, c -
9.469(12) A, da = 1.176 at 126 kbar and
7.959(5) A, 9.353(17) A, 1.175 respec-
tively at 144 kbar. Since the intensities
from the 8* - phase are large both in the x-
ray diffraction and the Raman scattering
experiments, this phase must be a signifi-
cant portion of the volume. From mass-
balance considerations, the composition of
the e* - phase is clearly larger than 5 mol %
He.

The experiment on e* with 1 0 % He was
performed at 92 kbar. Before the experi-
ment was done, the sequence of transitions
leading to this phase was verified, as well
as the vibrational Raman spectrum. The
sample consisted of several crystals, but no
additional powder, probably due to the
presence of a small amount of the helium-
rich fluid phase. Strong reflections includ-
ing series of overtones were obtained (see
Fig. 79). From the measured d-spacings, a
hexagonal unit cell with a - 8.258(2) A, c

140

CARNEGIE INSTITUTION

100000

' t t t I

■

Energy (keV)

Fig. 79. Diffraction pattern of e* with 10 % helium
taken at 26=9°. The fundamental reflection of the
(101) class at 13.7 keV and 4 overtones at 27.3,
41.1, 54.7 and 68.4 keV are indicated by the
arrows. The peaks near 18 keV are escape peaks
from the first overtone and the other peaks origi-
nate from the gasket.

= 9.747(5) A and c/a=1.180 was obtained,
similar to the results at 5 mol% He. Four
reflections could be attributed to one crys-
tal and five reflections to a second one.

The reflections of the e* - phase do not
fulfill the condition -h+k+l=3n of rhombo-
hedral structures. Therefore, we can rule

out the R3c space group for this phase.
However, the contents of the unit cell can
be estimated on the basis of the available
information. We assume the phase to be
stoichiometric and calculate the free en-
thalpy difference between this phase and
the coexisting phases, using known p(V)
isotherms of N2 (Olijnyk, 1 990) and He (Le
Toullec et ai, 1990) and making assump-
tions for the volumes of the mixture that are
justified by theory and simulations. It then
turns out that the e* - phase becomes stable
when it contains at least 22 N2 molecules
and 2 He atoms per unit cell. This also
yields a volume difference with the coex-

isting phases that is reasonable in view of
the pressure jumps measured previously at
transitions involving this phase (Vos and
Schouten, 1990; Vos, 1991). Furthermore,
this is also consistent with the fact that the
Raman shifts of this phase are very similar
to those of the 8 and e - phases at the same
pressure (Scheerboom and Schouten, 1 99 1 ).
Computer simulations were performed
using a variable shape simulation cell
(Rahman-Parinello method) to allow for
crystal transformations. The best available
potentials for N2 and He were used (for
details see de Kuijper, 1991). At 300 K and
pressures of 200 and 300 kbar, it was found

that pure N2 remains in the R3c space
group. This result was obtained with both
hexagonal and rhombohedral simulation
cells, which would have permitted the ap-
pearance of the alternative structures that
were observed earlier (Nose and Klein,
1986; Belak et al, 1990) and therefore
agrees with the experiment. For compari-
son with pure N2, simulations were per-
formed on N2-He, starting from the R3c
space group. It turns out that this structure
is maintained, indicating that it is at least
metastable for the mixture. An interesting
result is that the da drops considerably,
from 1.30 in pure N2 to 1.17 at 10 mol %
He.

From the present x-ray experiments, it
can be concluded that at room temperature
there is a negligible solubility of He in the
8 - phase of N2. Furthermore, the solubility
in the e* - phase is clearly larger than 5 mol
%, in agreement with previous experiments.
However, it turns out that He stabilizes a
phase with a different structure than pure e-

GEOPHYSICAL LABORATORY

141

N2. This phase has a smaller c/a ratio than
pure N2 at the same pressure, which is
consistent with computer simulations.

If the e* - phase turns out to be stoichio-
metric, the behavior of N2-He documented
here bears some similarity with that found
for colloidal suspensions. There, it was
recently found that mixtures with a size
ratio of 0.61 (cf., roughly 0.6 for helium
and nitrogen) form a stoichiometry with a
structure that is not encountered in the pure
components (Bartlett et al., 1990).

Scheerboom, M. I. M., and J. A. Schouten, Detec-
tion of the e-5 phase transition in N2 and the
N2-He mixture by Raman spectroscopy: new
evidence for the solubility of fluid He in solid
N2,/. Phys.: Condens. Matter, in press, 1991.

Vos, W. L., Phase equilibria in simple systems at
high pressure, Ph. D. dissertation, Universiteit
van Amsterdam, 1991.

Vos, W. L., and J. A. Schouten, Solubility of fluid
helium in solid nitrogen at high pressure, Phys.
Rev. Lett., 64, 898-901, 1990.

Evidence for Orientational Ordering of
Solid Deuterium at High Pressures*

Russell J. Hemley and Ho-Kwang Mao

References

Bartlett, P., R. H. Ottewill, and P. N. Pusey,
Freezing of binary mixtures of colloidal hard
spheres,/. Chem. Phys. ,93, 1299-1312, 1990.

Belak, J., R. LeSar, and R. D. Etters, Calculated
thermodynamic properties and phase transi-
tions of solid N2 at temperatures 0<T<300 K
and pressures 0<p<100 GPa, /. Chem. Phys.,
92,5430-5441,1990.

Cromer, D. T., R. L. Mills, D. Schiferl, and L. A.
Schwalbe, The structure of N2 at 49 kbar and
299 K, Acta Cryst. B37, 8-11, 1981.

de Kuijper, Jhr. A., Computer simulations of phase
equilibria in molecular systems, Ph. D. disser-
tation, Universiteit van Amsterdam 1991.

Le Toullec, R., P. Loubeyre, and J. - P.Pinceaux,
Refractive-index measurements of dense he-
lium up to 16 GPa at T=298K: Analysis of its
thermodynamic and electronic properties,
Phys. Rev. B40, 2368-2378, 1990.

Mao, H. K., and P. M. Bell, Design and operation
of a diamond-window, high-pressure cell for
the study of single-crystal samples loaded
cryogenically, Carnegie Instn. Washington
Year Book, 79, 409-411, 1980.

Mills, R. L., D. T. Cromer, B. dinger, Structures
and phase diagrams of N2 and CO to 1 3 GPa by
x-ray diffraction, /. Chem. Phys., 84, 2837-
2845, 1986.

Nose, S., and M. L. Klein, Constant-temperature -
constant-pressure molecular-dynamics calcu-
lations for molecular solids: Application to
solid nitrogen at high pressure, Phys. Rev.
B33, 339-342, 1986.

Olijnyk, H., High-pressure x-ray diffraction stud-
ies on solid N2 up to 43.9 GPa, /. Chem. Phys.,
93, 8968-8972, 1990.

At low pressures hydrogen forms an
insulating molecular solid with the mol-
ecules in states of complete rotational dis-
order over a wide range of temperature.
With increasing pressure, the rotational
motion of the molecules is expected to
become more restricted, ultimately leading
to orientational ordering. Detailing the na-
ture of possible ordering transitions is im-
portant for understanding the mechanism
of pressure-induced metallization, a transi-
tion with important implications for both
condensed-matter and planetary physics.
Raman measurements of the high-fre-
quency intramolecular vibrational mode
(vibron) indicate that the molecular solid
remains stable to at least -250 GPa but
undergoes a phase transition at 150 GPa (at
77 K; Hemley and Mao, 1988). The low-
frequency vibrational spectrum provides
information on the state of ordering in the
solid and constraints on the crystal struc-

* This work was supported by NSF (DMR-89
12226 and EAR-8904080) and NASA (NAGW-
1722).

142

CARNEGIE INSTITUTION

ture at these pressures, which are beyond
the range of current x-ray diffraction tech-
niques (Mao etal., 1988). Previously, we
reported measurements of the evolution of
the low-frequency rotational bands and lat-
tice phonon of hydrogen to 162 GPa at 77-
295 K (Hemley et al.y 1990a). Over this
pressure interval the rotational bands per-
sist but broaden and the lattice phonon,
which correlates with the Z?2g optical pho-
non of the hexagonal-close packed struc-
ture, shifts continuously. The continuity of
the low-frequency bands as a function of
pressure indicates that an underlying hex-
agonal structure persists into the high-pres-
sure phase above 1 50 GPa .

Examination of the low-frequency spec-
trum of deuterium is important for under-
standing isotope effects on a variety of
properties of hydrogen at high densities.
Orientational ordering is energetically fa-
vored in the heavier isotope as a result of its
smaller rotational constant (#D2 = 29.9
cm-1 versus #H2 = 59.3 cm-1 in the gas
phase), which results in a stronger mixing
of free molecule rotational states in con-
densed phase. At low densities, changes in
temperature result in variations in the rela-
tive population of ortho and para species
(even and odd / rotational states). How-
ever, at high densities the single molecule
ortho-para distinction breaks down as a
result of mixing of rotational states, and
this is expected to occur at lower densities
in D2. Evidence for this is found in the
differences in the pressure at which sym-
metry breaking occurs in the 7=0 solids at
very low temperatures (Silvera and
Wijngaarden, 1981). Isotope effects are
also observed in the pressure dependence

Hydrogen
93.8 GPa

200 400 600 800 1000

Wavenumber, cm"1

1200

Fig. 80. Examples of low-frequency Raman spec-
tra of hydrogen and deuterium at 77 K. The two
broadened low-frequency bands observed in H2
correlate with the So(0) and S\(0) rotational tran-
sitions. The intense band observed in D2 at 250
cm-1 is identified as a libra tional mode in an
orientationally ordered structure. The optical pho-
nons observed in both isotopes are indicated.

of the molecular vibron; the shift is signifi-
cantly stronger in hydrogen relative to deu-
terium (see Hemley et al., 1991). It is
therefore of interest to determine whether
or not this difference is associated with
structural differences between the two iso-
topes. Finally, a distinct isotope effect is
observed in the pressure of the low-tem-
perature 150-GPa phase transition, which
is >10 GPa higher in deuterium. Under-
standing the origin of this effect is of inter-
est because the transition appears to be
associated with changes in electronic prop-

GEOPHYSICAL LABORATORY

143

erties, such as metallization (Hemley etal.,
1990a; Hemley and Mao, 1990).

Samples were loaded at room tempera-
ture in a modified Mao-Bell diamond-cell
with composite rhenium/T301 -stainless
steel gaskets. Here we report the results of
four separate experiments on deuterium
carried out at pressures from 20 GPa to
above 100 GPa and temperatures between
77 K and 295 K. Raman spectra were
measured using optical techniques de-
scribed previously (Hemley and Mao, 1988;
Hemley et aL, 1990a). Low-frequency
Raman spectra show a strong, weakly pres-
sure dependent, band at 240 cm-1, together
with a weaker peak at higher frequency
which exhibits a large pressure dependence.
A representative low-frequency Raman
spectrum of deuterium at 99 GPa is com-
pared with that measured for hydrogen at
similar pressures in Fig. 80. On the basis of
x-ray diffraction measurements (Mao et
aL, 1988; Hemley et aL, 1990b) and the
continuity of the spectra with increasing
pressure, the higher frequency band is iden-
tified as the E2g Raman-active phonon in
the hexagonal-close packed structure,
analogous to that found for hydrogen.

The low-frequency spectrum of hydro-
gen at 77 K below 100 GPa is characterized
by two broadened rotational bands [So(0)
and Si(0), corresponding to the AJ = 2, / =
0-2 and AJ = 2, /= 1 -3 excitations in the free
molecule]. The persistence of these bands
to 100 GPa was interpreted as an indication
of free (or nearly free) rotation of the mol-
ecules. In contrast, the low-frequency band
observed in D2 does not fit a rotational
transition: i.e., the frequency of the band is
240 cm~l, whereas the frequencies of the

So(0) and Si(0) occur at 6£ = 180 cm"1 and
10# = 300 cm-1. We interpret this band as
indicative of a new high-pressure molecu-
lar phase of deuterium. In view of the
presence of the optical phonon, we suggest
that the phase is an ordered (or partially
ordered) form with a hexagonal close-
packed structure and that the low-frequency
band is associated with librational motion.
The latter is consistent with its weak pres-
sure dependence.

Further evidence for ordering is found
in the frequency shift of the optical phonon.
The volume dependence of the optical pho-
non frequency for the two isotopes is shown
in Fig. 81. The D2 curve is parallel to that
obtained by Wijngaarden et aL (1983) at

E
o

i_r

PHONON
77 K

200 _

Volume (cm3/mol)

Fig. 81. Volume dependence of the optical pho-
non for H2 and D2: squares, present work; circles,
Wijngaarden etal. (1983). The volume was calcu-
lated from the pressure using the experimental
equation of state determined to 26.5 GPa at room
temperature (Mao et aL, 1988; Hemley et aL,
1990). The dotted line shows the frequencies
expected for D2 on the basis of the measurements
forH2. The data of Wijngaarden etal. (1983) have
been corrected using the new equation of state.

144

CARNEGIE INSTITUTION

much lower pressures, as noted previously
for H2 (Hemley et al, 1990a). At low
compressions, the frequencies of the modes
differ by a factor of V2 as a result of the
differences in masses [i.e., (mnz/^D2)^],
as expected if the modes (assumed to be
harmonic), crystal structure, and volume
are identical for the two isotopes. It is
evident, however, that this relationship does
not hold at higher compressions: the D2
curve is higher than that expected on the
basis of the measurements for H2. This
offset may arise from a modification of the
crystal structure by ordering, which could
affect the frequency of the optical phonon
owing to changes in intermolecular inter-
actions in the ordered state (even at con-
stant volume). Alternatively, we note that
orientational ordering should result in a
more efficient packing of the molecules
relative to the rotationally disordered state,
and that the frequency of the phonon is a
strong function of volume. Thus, a second
possibility is that the offset indicates that
the molar volume of deuterium (which is
ordered) is lower than that of hydrogen
(which appears not to be fully ordered) at
the same pressures. This needs to be exam-
ined by low-temperature x-ray diffraction.
The available diffraction data provide some
evidence for a lower volume in D2 even at
room temperature (-2% at 30 GPa), so it is
possible that effects of rotational ordering
are present at room temperature (Fig. 81).
It should be pointed out that if the volume
difference at 77 K persists to higher pres-
sure (>150 GPa), it may contribute to the
isotope effect on the pressure of the high-
pressure phase transition (i.e., the low-
pressure phase is stabilized to higher pres-

sures in the heavier isotope).

The results may be compared with the
measurements of Silvera and Wijngaarden
(1981), who studied ortho-D2 at 5 K to 54
GPa. They reported the observation of a
broadened low-frequency band at 220-240
cm- * above 28 GPa, which they interpreted
as arising from an ordering-type transition
(broken symmetry transition in the J = 0
molecules). Since Silvera and Wijngaarden
(1981) were unable to measure the optical
phonon above the transition, they proposed
that the high-pressure phase has the cubic
Pa3 structure. The present observations of
the optical phonon indicate that the struc-
ture of the solid at 77 K (and above) is not
Pa3 because the Raman-active excitations
in this structure comprise only librational
modes (phonon is inactive). The close simi-
larity in the librational modes measured in
the two studies strongly suggests that the
transition observed at 5 K also takes place
within the hep-type structure. This assign-
ment is consistent with the results of recent
theoretical calculations which indicate that
structures based on hep are stable relative
to cubic (e.g., Pa3) at high densities (Raynor,
1987; Barbee etai, 1989; Ashcroft, 1991;
Kaxirase/a/., 1991).

Kaxiras et al. ( 1 99 1 ) have performed an
extensive series of calculations of different
molecular ordering schemes within hep. A
new class of oriented hexagonal structures
based on a herring-bone type configuration
has been found to be energetically favored
and to have larger band gaps than that of the
structure assumed in previous work. We
propose that the structure of the phase of D2
observed here is closely related to these
structures. This assignment is consistent

GEOPHYSICAL LABORATORY

145

Deuterium

130GPa

77 K

138 GPa
295 K

200 400 600 800 1000

Wavenumber, cm-1

1200

Fig. 82. Raman spectrum of deuterium at 130-138
GPa at 77 K and 295 K.

with the experimental evidence for a band
gap (insulating state) to high pressures
(i.e., to at least -150 GPa), which is one of
the key problems with previously proposed
ordered hexagonal structures [see, Ashcroft
(1991) and Kaxiras etal. (1991)]. Confir-
mation of this structure should be possible
by low-temperature single-crystal x-ray
diffraction. Further work is also required to
determine the P-T stability field of the
phase as function of pressure, temperature,
and ortho-para state (at lower pressures).
The librational band weakens gradually
with increasing pressure above 100 GPa,
and diamond fluorescence tends to increase
at these pressures. As a result, the phonon
could not be measured above 100 GPa,
although the stronger low-frequency band
is readily apparent (Fig. 82). With increas-
ing pressures above -140 GPa at 77 K the

band appears to weaken somewhat but was
observed through the high-pressure phase
transition at -165 GPa, despite the marked
discontinuity in the vibron frequency
(Hemley etal., 1991). Hence, D2 is appar-
ently ordered over the entire pressure inter-
val of this study at 77 K, although the
broadening of the band at room tempera-
ture (Fig. 82) may indicate that the mol-
ecules are disordered at higher tempera-
tures. At still higher pressures, an increase
in scattering intensity is observed in the
vicinity of the low-frequency band, a de-
tailed study of which will be presented
elsewhere.

References

Ashcroft, N. W., Optical response near a band
overlap: Application to dense hydrogen, in
Molecular Systems under High Pressure, R.
Pucci and G. Piccitto, eds., pp. 201-222,
Elsevier, Amsterdam, 1991.

Barbee, T. W., A. Garcia, M. L. Cohen, and J. L.
Martins, Theory of high-pressure phases of
hydrogen, Phys. Rev. Lett. 62, 1150-1153,
1990.

Hemley, R. J., and H. K. Mao, Phase transition in
solid molecular hydrogen at ultrahigh pres-
sures, Phys. Rev. Lett., 61, 857-860, 1988.

Hemley , R. J., andH. K. Mao, Critical behavior in
the hydrogen insulator-metal transition, Sci-
ence, 249, 391-393, 1990.

Hemley, R. J., H. K. Mao, and J. F. Shu, Low-
frequency vibrational dynamics and structure
of hydrogen at megabar pressures, Phys. Rev.
Lett. 65, 2670-2673, 1990a.

Hemley, R. J., H. K. Mao, L. W. Finger, A. P.
Jephcoat, R. M. Hazen, and C. S. Zha, Equa-
tions of state of solid hydrogen and deuterium
from single-crystal x-ray diffraction to 26.5
GPa, Phys. Rev. B 42, 6458-6470, 1990b.

Hemley, R. J., H. K. Mao, and M. Hanfland,
Spectroscopic investigations of the insulator-
metal transition in solid hydrogen, in Molecu-
lar Systems under High Pressure, R. Pucci
and G. Piccitto, eds., pp. 223-243, Elsevier,
Amsterdam, 1991.

Kaxiras, E., J. Broughton, and R. J. Hemley, Onset
of metallization and related transitions in solid

146

CARNEGIE INSTITUTION

hydrogen, Phys. Rev. Lett., 67, 1138-1141,
1991.

Mao, H. K., A. P. Jephcoat, R. J. Hemley, L. W.
Finger, C. S. Zha, R. M. Hazen, andD. E. Cox,
Synchrotron x-ray diffraction measurements
of single crystal hydrogen to 26.5 GPa, Sci-
ence 239, 1131-1134,1988.

Raynor, S., Novel ab initio self-consistent-field
approach to molecular solids under pressure.
II. Solid H2 under high pressure, /. Chem.
Phys., 87, 2795-2799, 1987.

Silvera, I. F., and R. J. Wijngaarden, New low-
temperature phase of molecular deuterium at
ultrahigh pressure, Phys. Rev. Lett., 47, 39-42,
1981.

Wijngaarden, R. J., V. V. Goldman, and I. F.
Silvera, Pressure dependence of the optical
phonon in solid hydrogen and deuterium up to
230 kbar, Phys. Rev., B 27, 5084-5087, 1983.

GEOPHYSICAL LABORATORY

147

BlOGEOCHEMISTRY

Nitrogen Isotope Tracers of
Atmospheric Deposition in Coastal
Shelf Waters off North Carolina.

Marilyn L. Fogel and Hans W. Paerl *

Nitrogen plays a key role in regulating
marine primary and secondary production
both on regional and global scales (Ryther
and Dunstan, 1971; Nixon et al, 1986).
Nitrogen sources in aquatic ecosystems
may be external ("new") or internal ("re-
cycled"). The relative utilization of "new"
and "recycled" N inputs by phytoplankton
is important in determining levels of pri-
mary and secondary production in coastal
ecosystems. The need for a more detailed
understanding of the dynamics of new vs.
recycled production in coastal waters is
pressing, because previously pristine seg-
ments of the coastal oceans are now exhib-
iting both incipient and advanced stages of
eutrophication (Cosper etaL, 1987; Paerl,
1988). Terrigenous point and nonpoint
inputs have traditionally been identified as
the most likely nutrient sources supporting
new production in heavily impacted, eutro-
phic estuaries, including the Chesapeake,
San Francisco, Delaware, and Narraganset
Bays (Boynton et al., 1982). The connec-
tion between man's watershed activities
and coastal eutrophication, however, ap-

Institute of Marine Sciences, University of North
Carolina

pears more cryptic in many other places
(e.g., South Atlantic Bight). Recently, ni-
trogen deposition from rainfall has been
shown to stimulate primary production in
coastal waters adjacent to North Carolina
(Paerl, 1985; Paerl et a/., 1990).

Atmospheric deposition, as wet and
dryfall, is an increasingly important source
of biologically-usable nitrogen in estua-
rine and coastal regions (Paerl, 1985;
Legendre and Gosselin, 1989). Large East
Coast estuaries and certain European seas
currently receive about 20-50% of their
combined nitrogen loading from atmo-
spheric sources (Fisher etaL, 1988; Prado-
Fiedler, 1990; Loye-Pilot et al, 1990).
Nitrogen from atmospheric deposition (AD)
may be a unique source in coastal waters, as
direct surface water deposition may occur
downstream of estuarine zones where much
of the terrigenous nitrogen has been as-
similated. Proper identification of differ-
ent N sources and their fluxes are of prime
concern in understanding and ultimately
controlling recent eutrophication problems
in coastal ecosystems.

We have tested the possibility that AD
represents both a unique and
biogeochemically significant source of ni-
trogen supporting new production in North
Carolina's coastal Atlantic waters by trac-
ing the fate of AD products into natural
phytoplankton assemblages through the use
of stable N isotope signatures. Nitrogen
isotopes at the natural abundance level
have been used extensively in tracing ei-
ther biochemical processes or sources of

148

CARNEGIE INSTITUTION

food in complex ecosystems (Owens, 1 987).
A study by Showers et al. (1990) in the
Neuse River of North Carolina demon-
strated distinct nitrogen sources on the ba-
sis of isotopic composition. The isotope
ratio of the nitrate from sewage treatment
or point sources was isotopically heavy
((515N =~12%o). Showers et al (1990)
were able to distinguish a source of nitrate,
enriched in 15N due to intense agricultural
activity, coming from nonpoint soil runoff
(<515N =7%o).

The study site for this investigation
was Bogue Sound and coastal waters di-
rectly off Morehead City and Beaufort,
North Carolina. In this region N sources
from sewage treatment are almost nonex-
istent. Therefore, remineralization and ag-
ricultural runoff provide the primary sources
of nitrogen to the phytoplankton in this
ecosystem. We also expected that the <515N
particulate nitrogen would be influenced
by recycled nitrogen, as 15N-enriched am-
monium (Cifuentes et al. , 1 989), and a 1 5N-
enriched nitrate from agricultural runoff,
providing these were the only two sources
of nitrogen for primary production.

Nitrogen in wet or dry deposition is
considerably more depleted in 15N relative
to recycled or agricultural inputs (Heaton,
1986). Variability in the isotopic composi-
tion of the nitrogen pools in acid deposition
may be indicative of certain atmospheric N
sources. For example, in industrial zones of

Europe, the <515N of dissolved ammonium
has a mean value of -12 %c, whereas an
average value for dissolved nitrate is -3 %c
(Freyer, 1979). Atmospheric nitrate as
NOx from industrial and automobile pollu-
tion should have a 615N near that of air (0

%c). If, however, the source of NOx is from
nitrification in soils, then owing to large
isotope fractionation (Mariotti et al. ,1981),

the 8 N of dissolved nitrate may be more
negative.

The concentration and isotopic compo-
sition of N in certain rainfall events was
determined. Collections of atmospheric
wet and dry deposition were made on the
roof of the Institute of Marine Sciences, a
location free of potential sources of con-
tamination (e.g., trees, powerlines, other
buildings). Large polypropylene pans hav-
ing splashproof walls were carefully acid-
washed (0.01 N HC1), then rinsed three
times with deionized water in order to
remove any traces of nutrients. Pans were
placed in elevated stands on the IMS roof-
top, prior to precipitation events, to collect
rainwater. Collectors were deployed just
prior to, and removed immediately after,
precipitation events in order to minimize
contamination.

Concentrations of dissolved inorganic
nitrogen species were first analyzed at the
University of North Carolina on a subsample
using the methods of Strickland and Par-
sons (1972). Rainwater samples were
shipped frozen to the Geophysical Labora-
tory and thawed just before isotope analy-
sis. Initial trials with rainwater indicate
that molecular sieve Zeolite W-85 adsorbed
the ammonium from the rainwater directly
without distillation. The <515N of solutions
containing a known NH + was within the
standard error of the measurement (± 0.5
%o) (Velinsky et al, 1989). After NH4+
removal, some aliquots were freeze-dried.
The residual material, containing the dis-
solved nitrate and any dissolved organic

GEOPHYSICAL LABORATORY

149

Atlantic Ocean

Cape

ookogt

65 km

Offshore Sampling Site

Fig. 83. Map of coastal North Carolina showing three sampling locations: Bogue Sound, Nearshore, and
Offshore sites.

150

CARNEGIE INSTITUTION

matter, was combusted for isotopic analy-
sis.

In bioassays and natural waters, the N
isotopic ratio of the whole phytoplankton
sample was measured. Samples were col-
lected from three locations: a coastal
nearshore site, Bogue Sound (near Beau-
fort Inlet), and 90 km offshore (Fig. 83).
These sites represent N-depleted, full-sa-
linity Atlantic coastal and mesohaline es-
tuarine waters. Hydrologically, Bogue
Sound is a portion of the meso- to euhaline
component of the Newport River Estuary.
During incoming tides, however, Bogue
Sound is a conduit for nearshore-coastal
Atlantic Ocean water. Chronic N limita-
tion characterizes both estuarine and coastal
waters transitting Bogue Sound (Thayer,
1978; Paerl, 1988). Water samples were
filtered first through Nitex screening to
remove zooplankton. Following initial
screening, remaining phytoplankton were
collected on Whatman GF/F filters (0.7 m

Z
in

NH4

NOx + DON

■

V A

-5-

10-

15 i
(

i
) 60

i
120

■

180

■ i • i

240 30(

■ i
) 360

Julian Days

Fig. 84. Nitrogen isotopic composition of NH4+
and NOx plus dissolved organic N (DON) in
continental rain events occurring on the North
Carolina coast over a year's time from 1 January
to 31 December (Julian Days). Samples were
collected from April-May 1988 (n=2) and August
1990- April 1991 (n = 6). Variations in <515N most
likely reflect differences in the sources of com-
bined N to atmosphere (e.g., agricultural input vs.
industrial input).

nominal pore size). Particulate organic
matter on glass fiber filters was ground
with CuO and placed in preheated quartz
tubes with copper metal. The quartz tubes
were evacuated and sealed off, combusted
batchwise at 910° C for 2 h, and cooled at a
controlled rate. Replicate analysis of fil-
ters plus organic material gave values with
standard deviations of ± 0.5 %o.

The <515N of the NH.+ from continental

rainfalls varied throughout the year (Fig.
84), but had values considerably more
negative than either recycled or agricul-
tural inputs. When nitrogen is incorporated
into phytoplankton during primary pro-
duction, the isotopic compositions of algae
will depend on (1) any biochemical frac-
tionations that may occur and (2) the iso-
tope ratios of the available nitrogen sources
(Cifuentes et al, 1989). Accordingly, the
<515N of primary producers should shift
with the addition of nitrogen from AD. In
fact, when significant rainfall events oc-
curred, the <515N of particulate nitrogen
decreased within a few days after the event
(Fig. 85).

•z.

15-

10-

A U

t *

D Offshore

/k Bogue Sound

■ Nearshore

0 30 60 90 120 150 180 210 240 270
Fall Winter Spring

Time (Days)

Fig. 85. Nitrogen isotopic composition of particu-
late material sampled from three coastal and estua-
rine sites off North Carolina as a function of time.
Day 1 = 1 August 1990. Arrows indicate the
timing of a continental rainfall event.

GEOPHYSICAL LABORATORY

151

^j"

bU -
50-
40-
30-
20-
10-
0-

•

Mixing depth = 1 .5 m

•

[PN] ( ji gN/L)

V • Am •* •

) 30

i
60

120

150 180 210 240

Time (Days)

30-

Q 20

•" 10

Mixing Depth= 2 m

[PN] (W N/L)

-i ■ r

V I tSafL

30 60 90 120 150 180 210 240

Time (Days)

Mixing Depths
■ 0.5m
• 5m

90 120 150 180 210 240
Time (Days)

mm Continental V~Z\ Mixed K/sA Oceanic

Fig. 86. Total dissolved inorganic nitrogen (DIN)
from atmospheric deposition that is mixed into
surface layers of the North Carolina Coast. All
rainfall events occurred from 1 August 1990 to
April 1991 (Table 22). (A) Bogue Sound site. (B)
Nearshore site. (C) Offshore site. Mixing depths
are approximations estimated from the depth of
the water column and the photic zone. The line
indicates average ambient particulate N concen-
trations measured periodically throughout the time
period. Levels of DIN were generally < 1 |imole
N/L.

Table 22. Dissolved inorganic N (DIN) content of selected significant (more than 0.5 cm)
rainfall events at the UNC Institute of Marine Sciences, Morehead City, North Carolina.*

Date

Origin [NOx]

[NH/]

Amount(cm)

9 Aug 90

3.58

Continental

659

1115.5

4.23

10 Sep 90

3.41

Continental

2213

1416.4

0.51

24-26 Oct 90

4.60

Mixed

295

103.4

3.98

6 Nov 90

5.74

Mixed

594.5

128

0.53

30 Nov 90

4.75

Mostly Oceanic

216.5

48.7

1.85

4 Dec 90

5.40

Oceanic

40.2

77.9

1.43

9 Dec 90

4.65

Mostly Oceanic

130.6

2.49

21 Dec 90

4.45

Mixed

280.4

74.5

1.19

22 Dec 90

4.81

Oceanic

60.1

9.13

0.77

3-4 Jan 91

4.17

Mostly Continental

434.4

84.8

0.44

9 Jan 91

4.79

Mostly Oceanic

75.2

46.8

4.66

12 Jan 91

4.79

Mixed

224

122

1.94

16 Jan 91

5.10

Oceanic

83.5

2.84

20 Jan 91

5.16

Oceanic

30.2

43.9

3.78

25 Jan 91

4.74

Mixed

118.8

28.9

2.27

8 Feb 91

4.89

Oceanic

44.6

1.72

5 Mar 91

5.13

Mostly Oceanic

87.2

76.2

5.30

14 Mar 91

4.36

Mostly Continental

376.7

275.6

1.56

30 Mar 91

5.07

Mostly Oceanic

159.9

187.3

2.24

* Weather fronts were documented from satellite imagery and dominant wind direction.
DIN concentrations are in Lig N/ L. ND indicates non-detectable concentrations.

152

CARNEGIE INSTITUTION

Table 23. Bioassay experiments to assay incorporation of atmospheric deposition into particulate
nitrogen from coastal samples.

Type of Date Final Rain Initial Final 5^N Rainwater Origin of Date of

Bioassay dilution S^N PN PN S^N NH44" last rain last rain

Cubitainer
Cubitainer
Mesocosm

13-Sep-90
17-Oct-90
27-Apr-91

1% +4.2

5% +10.8
1.80% +16.3

+4.4 ND* Continental 10-Sep-90

+3.5 -9.5 Mixed 22-Sep-90

+12.2 -1.5 Mixed 22-Apr-91

+15.9 ND Oceanic 27-Apr-91

Initial particulate nitrogen samples were collected from Bogue Sound, not determined.

Rainfall events in the Beaufort, North
Carolina, area originate from either conti-
nental or oceanic sources, or are frequently
a mixture of the two (Table 22).

The amount of dissolved inorganic ni-
trogen (DIN) and pH can usually be related
to the origin of the rainfall event. Continen-
tal storms had an average DIN concentra-
tion of 1643 jig N/L (range = 519-3629
jigN/L), whereas oceanic events contained
as little as 150 jug N/L (range = 69-347
jigN/L). To assess the relative importance
of AD nitrogen to the ecosystem, the flux of
N falling on the three coastal sites has been
calculated for each rainfall event. Depend-
ing on the mixing depth at each location
and the amount of rainfall, rain from all
three origins contributed significant
amounts of nitrogen, especially at the off-
shore site (Fig. 86).

To determine direct uptake of nitrogen
from AD, short and long-term bioassays
were designed to examine impacts of AD
on isotopic composition and primary pro-
duction under natural irradiance and tem-
perature conditions. Two independent bio-
assays were employed. The first used rela-
tively small volume Cubitainers (4L) incu-
bated in situ for delineating short-term (1
day-1 week) impacts (Paerl et a/., 1990).
The second used previously constructed

mesocosms (670 L), designed to evaluate
longer-term chronic loading effects and to
approximate the estuarine environments
surrounding the Beaufort-Morehead City
area. When a rainfall event was antici-
pated, a control set of mesocosms was
covered with transparent polyethylene,
thereby excluding rain water. A second set
of mesocosms remained uncovered, thereby
allowing rain input. To a third set of
mesocosms, rainfall from a previous con-
tinental event was added. Added rainfall
simulated dilutions commonly experienced
in coastal waters (Table 23).

An aliquot of 14C-NaHC03 was
added to each vessel in order to monitor
photosynthetic 14C02 assimilation as a
measure of microalgal production. A par-
allel 14C-free set of vessels were deployed
for stable isotope analyses. Initial samples
for particulate <515N analyses were taken at
this time. After 4 days, samples were col-
lected and filtered for 14C02 assimilation
and particulate <515N.

The first Cubitainer experiment fol-
lowed directly after a continental rainfall
that contained a significant amount of DIN
(Table 23). Although there was some stimu-
lation of primary production as measured

with 14C uptake, no isotopic changes re-
sulted. In the second experiment, the added
rainfall was from 9 Aug 1990 water, which

GEOPHYSICAL LABORATORY

153

had a 515N-NH,+ of -9.5 %o. Prior to this

experiment, no significant rainfall had oc-
curred in the area for 25 days. The addition
of 5 % rainwater to the Cubitainer caused a
7 %o decrease in the <515N of the particulate
N, indicative of the uptake of the isotopi-
cally-light N from the acid rain.

The mesocosm experiment was con-
ducted following a period when rain fall
events had been primarily of oceanic or
mixed origin (see Table 22). Rain from a
mixed event (22 Apr 91) was added to a 1 .8
% dilution to certain mesocosms, while
others received oceanic rainfall that oc-
curred a day later (Table 23). Primary
production was stimulated with the addi-
tion of the 22 April rainwater, and 5 N of
particulate N decreased by 4 %o. In con-
trast, no stimulation of primary production
as determined by 14C uptake was mea-
sured, relative to controls, with the oceanic
event, and no change in the 5 N of particu-
late N was detected.

Previous results suggest that atmo-
spheric N inputs, which may supply as
much as 20 to 50 % of coastal ocean "new"
nitrogen inputs, represent a significant
source of N, that contribute to current rates
of eutrophication. By demonstrating the
actual incorporation of nitrogen from AD
into primary producers with N isotope trac-
ers, we conclude that the biological mea-
surements of enhanced primary productiv-
ity are due to the presence of this alterna-
tive source of a limiting nutrient. Atpresent
we do not know how primary productivity
alterations in response to acid rain might
shape short- or long-term production trends
of estuarine and coastal ocean food chains.
This problem is by nature a global one. The

oxides of nitrogen responsible are largely
generated in major urban and industrial
centers located to the north and west of
North Carolina, although their exact origin
can only be speculated upon on the basis of
concentration and location. Fertilizers and
animal wastes associated with farming may
also be important sources of NH + in this
coastal area. Isotope tracers of atmospheric
deposition may not only be used in tracing
N into the food chain, but also may have
some relevance to the origin of the atmo-
spheric nitrogen.

References

Boynton, W. R., W. M. Kemp, and C. W. Keefe,
A comparative analysis of nutrients and other
factors influencing estuarine phytoplankton pro-
duction, in Estuarine Comparisons, V. S.
Kennedy, ed., pp. 69-90, Academic Press, NY,
1982.

Cifuentes, L. A., M. L. Fogel, J. R. Pennock, and
J. H. Sharp, Biogeochemical factors that influ-
ence the stable nitrogen isotopic ratio of dis-
solved ammonium in the Delaware estuary,
Geochim. Cosmochim. Acta, 53, 2713-2721,
1989.

Cosper, E. M, W. C. Dennison, E. J. Carpenter, V.
M. Bncelj, J. G. Mitchell, S. H. Kuenstner, D.
Colflesh, and M. Devey, Recurrent and persis-
tent brown tide blooms perturb coastal marine
ecosystem, Estuaries, 10, 284-290,. 1987.

Fisher, D., J. Ceraso, T. Mathew, and M.
Oppenheimer, Polluted Coastal Waters: The
Role of Acid Rain. Environmental Defense
Fund, New York, 1988.

Freyer, H. D, Seasonal trends of NH4+ and N03'
nitrogen isotope composition in rain collected
at Julich, Germany, Tellus, 30, 83-92, 1979.

Heaton, T.H.E., Isotopic studies of nitrogen pollu-
tion in the hydrosphere and atmosphere: A
review, Chem. Geol, 59, 87-102, 1986.

Legendre, L. O. and Gosselin, M. New production
and export of organic matter to the deep ocean:
consequences of some recent discoveries,
Limnol. Oceanogr., 34, 1374-1380, 1989.

Loye-PiJot, M. D., J.M. Martin, and J. Morelli,
Atmospheric input of inorganic nitrogen to the
western Mediterranean, Biogeochem., 9, 1 17-
134, 1990.

154

CARNEGIE INSTITUTION

Mariotti, A., J. C. Germon, P. Hubert, P. Kaiser, R.
Letolle, A. Tardieux, and P. Tardieux, Experi-
mental determination of nitrogen kinetic iso-
tope fractionation, some principles; illustration
for the denitrification and nitrification prin-
ciples, Plant and Soil, 62, 413-430, 1981.

Nixon, S. W., C. A. Oviatt, J. Frithsen, and B.
Sullivan, Nutrients and the productivity of es-
tuarine and coastal marine ecosystems. /.
Limnol. Soc. So. Afr., 12, 43-71, 1986.

Owens, N. J. P., Natural variations in 15N in the
marine environment, Adv. in Mar. Biol., 24,
411-451,1987.

Paerl, H. W., Enhancement of marine primary
production by nitrogen-enriched acid rain, Na-
ture, 316, 747-749, 1985.

Paerl, H. W., Nuisance phytoplankton blooms in
coastal, estuarine, and inland water, Limnol.
Oceanogr., 33, 823-847, 1988.

Paerl, H. W., J. Rudek, and M. A. Mallin, Stimula-
tion of phytoplankton production in coastal
waters by natural rainfall inputs: Nutritional
and trophic implications, Mar. Biol,. 107, 247-
254, 1990.

Prado-Fiedler, R., Atmospheric input of inorganic
nitrogen species to the Kiel Bight, Helgolander
Meersuut, 44, 21-30, 1990.

Ryther, J. H., and W. M. Dunstan, Nitrogen,
phosphorus, and eutrophication in the coastal
marine environment, Science, 171, 1008-1013,
1971.

Showers, W. J., D. M. Eisenstein, H. W. Paerl, and
J. Rudek, Stable isotope tracers of nitrogen
sources to the Neuse River, North Carolina,
Water Resources Institute Report No.253, 1990.

Strickland, J. D. H. and T. R. Parsons, A Practical
Handbook of Seawater Analysis. Fish. Res.
Board Can. Bull. 167, 1972.

Thayer, G. W., Identity and regulation of nutrients
limiting phytoplankton production in the shal-
low estuaries near Beaufort, N.C., Oecologia,
14, 75-92, 1978.

Velinsky, D. J., L. A. Cifuentes, J. R. Pennock, J.
H. Sharp, and M. L. Fogel, Determination of
the isotopic composition of ammonium-nitro-
gen at the natural abundance level from estua-
rine waters, Marine Chemistry, 26, 351-361,
1989.

Nitrogen Diagenesis in Anoxic Marine
Sediments: Isotope Effects

David J. Velinsky, David J. Burdige* and
Marilyn L. Fogel

Due to the complexity of particulate
nitrogen (PN) transformations, the stable
isotopic composition of sedimentary nitro-
gen has received little attention as an indi-
cator of changes in the nitrogen biogeo-
chemistry of the oceans. During diagenesis
nitrogen can change oxidation state (-3 to
+5) through a series of bacterially - medi-
ated reactions (Klump and Martens, 1983).
These transformations have a range of iso-
tope effects, related mainly to kinetic dif-
ferences in the reactivity between the light
(14N) and heavy (15N) isotopes (Wada et al. ,
1 975). As a result, diagenetic processes can
affect the overall isotopic composition of
dissolved and particulate nitrogen within
the sediments and overlying waters. For
example, if the extent of denitrification
(N03" — > N2) in the water column and
continental shelf sediments changed over
time, the #5N of oceanic nitrogen could
shift significantly due to the large isotope
effect associated with denitrification (e =
30-40 %c>; Wada et a/., 1975; Cline and
Kaplan, 1975) and the extent of this process
as a sink for combined nitrogen in the
oceans (Christensen, 1987). The source
and isotopic composition of nitrogen to the
sediments can change depending on con-
ditions in surface waters. Rau et al. (1987)

* Department of Oceanography, Old Dominon
University, Norfolk, Virginia

GEOPHYSICAL LABORATORY

155

showed that the #5N of PN in organic
carbon-rich Cretaceous marine sediments
were isotopically light (-4.0 to +1.0 %6)
compared to typical marine sediments (gen-
erally >4 %6). They speculated that N2 fixa-
tion was the dominant source of organic
nitrogen to these Cretaceous sediments.

The type and isotopic composition of
particulate nitrogen that is incorporated
into the sediments depends on source varia-
tions and isotope effects during diagenesis.
For source variations to be evaluated, it is
important to determine if any post-deposi-
tional isotope effect occurs during early
diagenesis. The purpose of this paper is to
investigate possible isotope effects during
diagenesis of PN in anoxic marine sedi-
ments. Once diagenetic effects are deter-
mined, a more accurate interpretation of
nitrogen isotopes in ancient and present^
day marine sediments can be obtained.

Framvaren Fjord

Great Marsh

▲ A

Leaves, Framvaren Fjord

-30

-25

-20

-15

513C

Fig. 87. The <513C of organic carbon and <515N of
particulate nitrogen from sediments taken from
the Chesapeake Bay, Framvaren Fjord and Great
Marsh. Also included are isotope values from
various plants around the Framvaren Fjord.

Methods and Study Areas

Concentrations and <515N of pore water
NH4+ and sedimentary PN were deter-
mined from cores taken in three contrasting
coastal marine environments: Framvaren
Fjord, Norway (FF; June 1989), Great
Marsh, Delaware (GM; June, 1988), and
Chesapeake Bay (CB; June, 1988).

The Framvaren Fjord is a permanently
anoxic fjord with concentrations of dis-
solved NH4+ and hydrogen sulfide (H2S)
in the bottom waters (maximum depth 180
m) of 2 mM NH4+ and 6 mM H2S, respec-
tively. FF sediments are permanently
anoxic. The sources of organic matter to the
sediments are primarily from bacterial and
phytoplankton production in the water col-
umn along with some terrestrial inputs (e.g.,
leaves).

The Great Marsh is a coastal salt marsh
dominated by the short form of Spartina
alterniflora. Marsh sediments in the upper
12 cm (i.e., the active root zone) cycle
seasonally betwenn oxidized and reducing
conditions (Velinsky and Cutter, 1991).
Sediments below -12 cm are permanently
anoxic. Sources of organic matter to these
sediments include Spartina production
and upland runoff.

Two sediment cores were obtained from
the Chesapeake Bay, near 38°56' N,
76°23'W, just south of the Chesapeake Bay
Bridge near Annapolis, Maryland. The
water column in this section of the bay
(maximum depth 30 m) is seasonally anoxic
or sub-oxic whereas the sediments are gen-
erally permanently anoxic (San Diego-
McGlone, 1991). Sources of organic mat-
ter to the sediments of this portion of the

156

CARNEGIE INSTITUTION

0
0

10 -

20 -

30 -

40 -

%PN and NH+ (mM)
1.0 2.0

515N

3.0

1 1 1 1 1 1 1—

—1 • 1 1

C/N ,

I \

\ nh; ■

PN A

i 1

_i i

10 15

C/N (atomic)

Fig. 88. The depth distributions of particulate nitrogen (PN), dissolved ammonium (NH4+) and the
organic carbon to particulate nitrogen ratio (C/N) along with the isotopic composition (515N) of
dissolved NH4+ and PN in Framvaren Fjord sediments.

bay are primarily derived from phy toplank-
ton production and to a lesser extent from
land runoff.

Box (CB) or gravity (FF and GM) cores
were sectioned at specific intervals. The
CB core was sectioned every 2 cm, and the
FF core was sectioned at 5-cm intervals.
The GM core was sectioned every 2.5 cm,
and due to the small volume of pore fluids
obtained and the low concentrations of
dissolved NH4+ in the upper 15 cm, the 2.5-
7.5 cm and the 10-15 cm sections were

combined for isotope analysis. Pore fluids
were separated from sediments by either
centrif ugation (FF) or with Reeburgh ( 1 967 )
sediment squeezers (CB and GM), then
filtered through Nuclepore 0.4 jiim filters.
Both sediment and pore waters were stored
frozen until sample preparation and analy-
sis.

The methods for the preparation and
determination of the <515N-NH4+and <515N-
PN are described in Velinsky et al. (1989)
and Cifuentes et al. (1988). The data are

GEOPHYSICAL LABORATORY

Table 24. Isotope discrimination between PN and NH4+.

157

Location

S15 PN*

Reference

Frarnvaren Fjord

3.3±0.9

-0.210.8

This Study

Chesapeake Bay

Core 41
Core 42

9.4±0.5
9.8±1.1

- 2.8 to - 0.7
2.3 to - 0.4

This Study
This Study

GreatMarsh,DE.

5.2±0.8

- 0.810.6

This Study

Santa Barabara Basin

SOG 005
GAS 24
GAS25
GAS 29

7.111.2
7.010.9
7.310.3
6.110.9

+ 2.611.5
+ 3.211.0
+ 4.110.6
+ 5.011.7

Sweeney and Kaplan (1980)

* Average ^l^PN

**Average A = #15NH4+ - 515PN

%PN and NH + (mM)
0 1.0 2.0 3.0

> p [

— I ' 1 r"

C/N ,

>v I /

\ NH4 "

PN 6

► >

) '

[

3 •

10 15 20

C/N (atomic)

Fig. 89. Similar to Fig. 88, except for the Great Marsh.

158

CARNEGIE INSTITUTION

reported in the standard S notation {i.e.,
<515N = [(#sample/#standard) - 1] x 103; where
R = 15n/14N} and the ratios are reported
against air (<515N = 0). Precision of repli-
cate samples for ammonium and PN isoto-
pic analysis is approximately ± 0.5 %o and
± 0.2 %o, respectively. Ammonium concen-
trations were determined with the proce-
dures described in Sharp et al. (1982) and
solid phase PN concentrations were deter-
mined with a Carlo-Erba ANA 1500 car-
bon and nitrogen analyzer.

Results and Discussion

To illustrate the differences between
sedimentary environments, the isotopic
compositions of organic carbon (<513C) and
particulate nitrogen (<515N) are plotted (Fig.
87). Because of the predominance of
Spartina aiterniflora (a C4 plant) in the
GM, the <513C of the sediments are greater
than those in either the FF and CB sedi-
ments. The carbon and nitrogen composi-
tion of FF sediments reflect terrestrial and

%PN
0.3 0.4

0.5

15.0

NH4 (mM)

Fig. 90a. Similar to Fig. 88, except for Chesapeake Bay, Core 41.

GEOPHYSICAL LABORATORY

159

E
o

Q.
CD

5.0

7.5

815N

10.0

12.5

15.0

NH^(mM)

Fig. 90b. Similar to Fig. 90a, except for Chesapeake Bay, Core 42.

water column-derived inputs, whereas the
CB sediments are isotopically enriched in
15N and 13C, indicating typical estuarine
phytoplankton.

In the FF sediments, the <515N of both
pore water NH4+ and sedimentary PN did
not change significantly with depth (Fig.
88). As such, the isotopic difference (A =
5!5N-NH4+ - <515N-PN) between the NH4+
and PN is small (~ -0.2%o) with the NH4+
being 15n depleted (Table 24). These data
indicate no expression of significant iso-
tope effect during ammonification of PN in
the FF sediments.

The concentrations of PN decreased
significantly (60%) in the GM sediments
(Fig. 89) from the surface to approximately
27 cm. Concurrently, the <515N-PN in-
creased very slightly in this interval with an
overall average of 5.66 ± 0.46 %o (n = 8).
Whereas pore water NH4+ concentrations
increased with depth, the <515N-NH4+ did
not change and averaged 4.30 ± 0.41 %o
(n=l). Similar to FF sediments, the <515N-
NH4+ was lighter than the <515N-PN, indi-
cating a small negative discrimination with
depth (i.e., 14N-NH4+ is preferentially re-
leased compared to 15N-NH4+; Table 24).

160

CARNEGIE INSTITUTION

This discrimination is a result of a number
of processes that could fractionate NH4+,
including ion-exchange, deamination of
amino acids (Macko and Estep, 1984), up-
take by bacteria within the sediments and
diffusion out of the sediments. The ob-
served distribution of <515N-NH4+ with
depth most likely results from a combina-
tion of these processes.

The Chesapeake Bay cores exhibited a
slightly different trend in <515N-NH4+ with
depth (Fig. 90a). In Core 41 , the concentra-
tion of PN decreased with depth in the

upper 10 cm and remains constant below
10 cm, while the <515N-PN did not change
in the upper 10 cm. Conversely, as the
concentration of NH44" increased with
depth, the <515N-NH4+ also increased. The
maximum increase in <515N-NH4+ was in
the same depth intervals as the decrease in
PN and increase in dissolved NH4+. The
differences between Core 41 and 42 (Fig.
90b) reflect the spatial heterogeneity of
sediments in the bay. A sharp maximum in
pore water NH4+ and <515N occurs in the 4-
6 cm interval and this was related to an

Nitrogen Cycle in Sediments

NO,

Water

uuj

2UUUUU

i«p

PN,

Burial

K-2.B

Sediment

no3

N
2

flux of PN to the sediments

denitrification

am monification

nitrate diffusion

k-2

microbial uptake

ammonium diffusion

nitrification

burial

Fig. 90. A conceptual model for the pathways of PN mineralization in marine sediments. Note that the
PN is broken into two fractions: a labile phase (PNl) and a refractory phase (PNr). With depth and burial,
the PNl fraction would decrease as would the rate of remineralization (k2; Burdige, 1991).

GEOPHYSICAL LABORATORY

161

extensive shell layer found at this depth.
This shell layer contained additional or-
ganic material that enhanced sulfate reduc-
tion causing the pore water sulfate concen-
trations to decrease to undetectable levels
at 4-6 cm versus 12-14 cm for Core 41. As
a result, pore water concentrations of dis-
solved NH4+ increased dramatically from
the surface (1.3 mM) to 4-6 cm (6 mM).
Below the NH4+ maximum, concentra-
tions decreased to near constant levels,
whereas the concentration of PN decreased
to a minimum at approximately 1 8 cm (Fig.
90b). The S^N of both NH4+ and PN
reflected the change in source of PN with
distinctly different isotopic ratios of nitro-
gen in the 4-6 cm interval.

The absence of any significant isotope
effect during PN remineralization may re-
flect the variety of processes that are affect-
ing PN in sediments. Each process could
have an intrinsic isotope effect that is fully
expressed in the overall solid phase PN
distribution. Alternately, the lack of any
observable shift in the <515N-PN with depth
may be due to a "mass" effect, in that only
a small fraction of the PN is remineralized
compared to the total PN. As such it would
be difficult to see any discrimination unless
it is very large. Pore water NH4+ should be
a better indicator of mineralized material
and, as such, should be a more accurate
reflection of any diagenetic isotope effect
on the PN.

The distributions of A are similar for the
Chesapeake Bay cores. In both cores, A in
the surface section was approximately -2.5
%o(i.e., the NH4+ is isotopically lighter than
the PN) and increased with depth. This
trend indicates a possible selective

remineralization in the upper sediments of
a more labile fraction of the PN with a
lighter isotopic composition (PNl; Fig.
91). This fraction of PN could represent
"fresh" phytoplankton material that has
reached the bottom sediments relatively
unaltered. Montoya et al. (1990) showed
that particulate material from the mainstem
of Chesapeake Bay during the spring had
nitrogen isotopic compositions of between
6.2 and 10.5 %o. If this scenario is the case,
the preferential degradation of this more
labile organic matter (Westrich and Berner,
1984; Burdige, 1991) would release dis-
solved NH4+ into the pore waters with
<515N similar to that of the source material
(Fig. 91). With depth less of this lighter
material remains and eventually the <5l5N-
NH4+ of the pore waters would reflect
remineralization of the "background", more
refractory, particulate nitrogen (PNr). Such
results were shown by Sweeney and Kaplan
(1980) for sediments taken from the Santa
Barbara Basin (Table 24). They demon-
strated that the <515N of the pore water
NH4+ reflected the degradation of a marine
source of PN with a <5i5N of approximately
10 %o. The range of <5i5N for total nitrogen
in these sediments was thought to be de-
rived from a mixture of a marine (i.e.,
phytoplankton at 10 %o) and terrestrial
sources (i.e., sewage-derived at 2 %c). Dis-
solved NH4+ in the pore waters was there-
fore postulated to be derived mainly from
the preferential degradation of the plank-
tonic organic nitrogen. Although the
data from Chesapeake Bay reflects the pos-
sible degradation of a labile fraction of PN,
the data from the Framvaren Fjord and
Great Marsh did not reflect this distribu-

162

CARNEGIE INSTITUTION

tion (Figs. 88 and 89; Table 24). It is pos-
sible that in the Framvaren Fjord the major-
ity of the degradation of PN occurred in the
water column above the sediments. By the
time this material reached the sediments
any labile phase was already degraded. In
the GM sediments, labile N could be re-
leased in the upper 10-12 cm and either
consumed by the Spartina or possibly
released via diffusion to the adjacent creek
waters

In conclusion, only a small isotopic
discrimination is expressed during diagen-
esis (Table 24). Whereas, the bulk Sl5N-
PN did not significantly change with depth,
it appears that selective remineralization of
a labile fraction of N may occur in certain
environments. This observation indicates
the <515N of specific fractions of the PN
could be used as a tracer of recently formed
organic material. Downcore variations in
<515N of solid-phase nitrogen probably re-
flect the material deposited to the sediment
surface. It is not possible to tell however, if
the isotopic composition of the PN formed
in the water column is reaching the sedi-
ment intact, only that the isotopic integrity
of the bulk PN appears to remain unaf-
fected.

References

Burdige, D. J., The kinetics of organic matter
mineralization in anoxic marine sediments,
Jour. Mar. Res., in press, 1991.

Christensen, J. P., J. W. Murray, A. H. Devol, and
L. A. Codispoti, Denitrification in continental
shelf sediments has major impact on the oce-
anic nitrogen budget, Global Biogeochem.
Cycles, 1(2), 97-116, 1987.

Cifuentes, L.A., J. H. Sharp, and M. L. Fogel,
Stable carbon and nitrogen isotope biogeo-

chemistry in the Delaware estuary, Limnol.
Oceangr.,33, 1102-1115, 1988.

Cline, J. D. and I. R. Kaplan, Isotopic fractionation
of dissolved nitrate during denitrification in
the eastern Tropical North Pacific, Mar. Chem.,
5,271-299, 1975.

Klump, J. V., and C. S. Martens, Benthic nitrogen
regeneration, in Nitrogen in the Marine Envi-
ronment, Chapter 12, EJ. Carpenter and D.G.
Capone, eds., Academic Press, New York, pp.
411-458,1983.

Macko, S. A., and M. L. F. Estep, Microbial
alteration of stable nitrogen and carbon isoto-
pic compositions of organic matter, Org.
Geochem., 6, 787-790, 1984.

Montoya, J. P., S. G. Horrigan, and J. J. McCarthy,
Natural abundance of 15N in particulate nitro-
gen and zooplankton in the Chesapeake Bay,
Mar. Ecol. Prog. Series, 65, 35-61, 1990.

Rau, G. H., M. A. Arthur, and W. E. Dean, 15N/14N
variations in Cretaceous Atlantic sedimentary
sequences: implication for past changes in
marine nitrogen biogeochemistry, Earth
Planet. Sci. Let., 82, 269-279, 1987.

Reeburgh, W. S., An improved interstitial water
sampler, Limnol. Oceanogr., 12, 223-234,
1967.

San Diego-McGlone, M. L., Processes affecting
the behavior of redox sensitive elements in
Chesapeake Bay, Ph.D dissertation, Old Do-
minion University, 1991.

Sharp, J. H., C. H. Culberson, and T. M. Church,
The chemistry of the Delaware estuary. Gen-
eral Considerations, Limnol. Oceanogr., 27,
1015-1028, 1982.

Sweeney, R. E. and I. R. Kaplan, Natural abun-
dances of 15N as a source indicator for near-
shore marine sedimentary and dissolved nitro-
gen, Mar. Chem., 9, 81-94, 1980.

Velinsky, D. J. and G. A. Cutter, Geochemistry of
selenium in a coastal salt marsh, Geochim.
Cosmochim.Acta,55, 179-191, 1991.

Velinsky, D. J., J. R. Pennock, J. H. Sharp, L. A.
Cifuentes and M. L. Fogel, Determination of
the isotopic composition of ammonium- nitro-
gen at the natural abundance level from estua-
rine waters, Mar. Chem., 26, 351-361, 1989.

Wada, E., T. Kadonaga and S. Matsuo, 15N abun-
dance in nitrogen of naturally occurring sub-
stances and global assessment of denitrifica-
tion from isotopic viewpoint, Geochem. Jour.,
9, 139-145, 1975.

Westrich, J. T., and R. A. Berner, The role of
sedimentary organic matter in bacterial sulfate
reduction: The G model tested, Limnol.
Oceangr., 29(2), 236-249, 1984.

GEOPHYSICAL LABORATORY

163

The Isotopic Ecology Of Plants And

Animals In Amboseli National Park,

Kenya

Paul L. Koch, Anna K. Behrensmeyer*
and Marilyn L. Fogel

Variations in the stable isotope ratios of
carbon ("C/1^), nitrogen ("N/^N), and
oxygen ("O/^O) provide information about
the ecology, physiology, and habitats of
living and extinct animals. For example,
the ^3C values of an animal's tissues are
controlled by the isotopic composition of
its diet, which, for herbivores, is related to
the photo synthetic pathway of food plants
(DeNiro and Epstein, 1978; Vogel, 1978).
Although affected by dietary #5N values,
N isotopes in animals vary with rainfall
amounts among ecosystems and among
trophic levels in an ecosystem (DeNiro and
Epstein, 1981; Heaton et ai, 1986).

We are investigating the isotopic
ecology of plants and animals in Amboseli
National Park, Kenya, for several reasons.
First, investigations of floral and faunal
isotopic composition in terrestrial ecosys-
tems are uncommon (e.g., Ambrose and
DeNiro, 1986; Sealy era/., 1987),andnone
evaluates C, N, and O isotopes simulta-
neously. Ecosystem studies test the gen-
erality of relationships determined either in
the laboratory or in comparisons of indi-
viduals from different regions. Second,
stable and radiogenic isotopes have been
employed to identify sources of elephant

*Department of Paleobiology, National Museum
of Natural History, Smithonian Institution, Wash-
ington, DC

ivory and rhinoceros horn, in order to con-
trol sales of poached versus legally hunted
animals (van der Merwe etal., 1990; Vogel
et ai, 1990). Ivory from various African
parks can be distinguished by its N, C, and
either Sr or Pb isotopic composition. How-
ever, if the isotopic composition of el-
ephants varies with time, because of habi-
tat, diet, or climate change, isotopic identi-
fication of source region may be unreli-
able. Either the isotopic composition of a
species must be constant through time
within an ecosystem, or the secular trends
must be minor when compared to differ-
ences between populations. Finally, isoto-
pic patterns in modern ecosystems can serve
as analogs for interpretation of the fossil
record. African faunas, with their diversity
of large mammals, are excellent analogs of
typical faunas before the Pleistocene ex-
tinction.

Study Area, Materials, and Methods

Amboseli Park is located in southern
Kenya (20°40'S,37°15'E; mean elevation,
1140 m). Annually, temperature averages
23°C and ranges from 15° to 31°C. Rain
falls in two seasons and averages 300 cm/
year. However, the park is continuously
supplied with spring water fed by melting
snow on Mt. Kilimanjoro. Habitats in the
park include grasslands, bushlands,
swamps, seasonal lakes, and woodlands.
Woodlands have retreated since the early
1970s, perhaps due to overbrowsing by
elephants or increased soil salinity. Tree
loss has altered the abundances of herbi-

164

CARNEGIE INSTITUTION

vores; there are more grazers (animals that
eat grass) and fewer browsers (animals
eating herbaceous and woody plants) and
mixed feeders (D. Western, pers. comm.).

Plant samples (mixtures of leaves and
stems) were collected in September 1 990 at
eight localities in the woodland, swamp,
swamp edge, plains, and bush habitats.
Faunal samples (tooth dentin or bone) were
collected from carcasses throughout the
park. Samples were collected from 1975
through 1990, and were in different states
of weathering. The minimum number of
years since death can be estimated from
weathering stage (Behrensmeyer, 1978).
For carcasses in advanced weathering
stages, however, determining actual time
since death is difficult.

Plants were air dried in the field, freeze-
dried in the laboratory, and then lightly
crushed. Bones and teeth were demineral-
ized with EDTA or 0.1 N HC1 to isolate
collagen (Tuross et ai, 1988), and then
treated with chloroform/methanol solution
to remove lipids. Plant and collagen samples
were placed in preheated quartz tubes with
CuO and Cu metal. Tubes were evacuated,
sealed, combusted at 910°C for 2 h, then
cooled at a controlled rate. Standard devia-
tions for analysis of standards were ±0.2%c
for #3C and S^N.

I so topic Variation in Amboseli Plants

The plants of Amboseli segregate into
two populations isotopically (Table 25, Fig.

92A). The 5 C values of grasses, which
use C4 photosynthesis, have a mean value

iu ■

aquatic

succulent

8 -

shrub & tree

□

herb

■

grass

.y e-

o 4-

- r

„

2 -

- X

n -

613C

8
tn
aJ

| 6

o 4

^ :,

□ elephant

□ mixed feeder

□ browser
■ grazer

Q carnivore

fiii

^ 00 O OJ ^t CD

613C

Fig. 92. (A) Histogram of carbon isotope compo-
sitions for Amboseli plants subdivided according
to physiogamy. (B) Histogram of carbon isotope
compositions for Amboseli mammals subdivided
according to feeding type.

± one standard deviation of -13.2 ± 0.9 %o.
The herbaceous plants and woody plants

employ CL photosynthesis and have mean
values of -27.1 ± 1.9 %o and -27.7 ± 2.4 V
respectively. This isotopic segregation

between CL shrubs, trees and herbs and C4
grasses is expected in a hot, dry region
(Tieszen and Boutton, 1988). Succulent
herbaceous and woody plants exhibit a

GEOPHYSICAL LABORATORY

165

Table 25. Isotopic data for Amboseli plants collected in 1990

Species

Family

Habitat

8^N

5l3C

Grasses and Sedges

Cynodon dactylon

Graminae (m)

swamp

10.6

-13.0

Sporobolus consimilis

Graminae (m)

swamp edge

8.9

-13.4

Sporobolus spicatus

Graminae (m)

swamp edge

8.6

-13.2

Sporobolus kentrophyllum

Graminae (m)

swamp edge

10.1

-13.4

Cynodon plectostachys

Graminae (m)

woodland

9.4

-14.2

Sporobolus helvolus

Graminae (m)

bush

8.8

-15.1

Sporobolus ioclades

Graminae (m)

bush

11.9

-13.0

Chloris roxburghiana

Graminae (m)

bush

8.7

-13.4

Chloris virgata

Graminae (m)

bush

10.6

-13.2

Enneapogon cenchroides

Gramiriae (m)

bush

9.4

-13.3

Cyperus immensus

Cyperaceae (m)

swamp

4.4

-11.2

Cyperus laevigatus

Cyperaceae (m)
Submerged Aquatic Plants

swamp

7.8

-12.2

Ceratophyllum sp. 1

Ceratophyllaceae

swamp

6.4

-23.0

Ceratophyllum sp. 2

Ceratophyllaceae
Herbs

swamp

9.3

-19.7

Pistia stratiotes

Araceae (m)

swamp

11.5

-29.0

Solanum incanum

Solanaceae

swamp edge

10.4

-25.0

Justicia odora

Acanthaceae

woodland

13.3

-27.6

Diplictera albicauda

n.d.

woodland

8.0

-27.8

Abutilon mauritanium

Malvaceae

plains

11.0

-29.7

Pluchea ovalis

Asteraceae

plains

8.1

-30.5

Cissampelos mucronata

Menispermaceae

plains

7.5

-27.7

Commicarpus sp.

Nyctaginaceae

plains

8.6

-28.9

Achyranthes aspera

Amaranthaceae

plains

11.4

-28.4

Withania somnifera

Solanaceae

plains

9.4

-25.2

Indigofera sp.

Leguminosae

bush

10.4

-25.0

Duosperma eremophiloum

Acanthaceae

bush

8.2

-27.1

Barleria spinisepala

Acanthaceae

bush

11.2

,24.3

Shrubs. Trees, and Succulent Plants

Trianthema ceratosepala

Aizoaceae

bush

12.8

-21.8

Sansevieria sp.

Agavaceae (m)

bush

12.6

-14.5

Euphorbia sp. 1

Euphorbiaceae

bush

13.3

-14.9

Euphorbia sp. 2

Euphorbiaceae

bush

13.9

-14.0

Sueda monoica

Chenopodiaceae

swamp edge

13.6

-13.3

Continued on next page

166

CARNEGIE INSTITUTION

Table 25. Continued

Species

Family

Habitat 515N <513C

Salvador a persica
Maerua triphyllum
Commiphora sp.
Boscia angustifolia
Acacia sp.
Azima tetracantha
Balanites glabra
Phoenix reclinata
Acacia tortilis

Shrubs and Trees

Salvadoraceae

swamp edge

5.3

-25.4

Capparidaceae

bush

8.8

-29.3

Burseraceae

bush

15.4

-29.3

Capparidaceae

bush

11.2

-24.6

Leguminosae

bush

10.2

-23.7

Salvadoraceae

woodland

7.6

-26.8

Balanitaceae

woodland

7.7

-29.9

Palmae (m)

woodland

7.1

-29.4

Leguminosae

woodland

8.9

-29.7

(m) indicates monocotyledons, all other plants are dicotyledons

range of 8 C values and may use either C.

or Crassulacean acid metabolism. Finally,

i ^
submerged aquatic plants have 5 C values

that can range between -12 and -33 %o,
depending on the pathway of carbon up-
take. Unlike terrestrial plants, which di-
rectly incorporate atmospheric CCL, sub-
merged plants may accumulate either dis-
solved C02 or HC03" (Raven, 1987).
Amboseli aquatic plants have #3C values
intermediate between C3 and C. plants.

The <515N values of Amboseli plants
form a unimodal distribution with a mean
of 9.8 ± 2.4 %o (Table 25, Fig. 93 A). This
mean value is slightly higher than that
reported by Sealy et al. (1987) for plants
from a region receiving 300 mm of rain.

Plant <515N values are not dependent on
either location or habitat type, although
plants from the bush habitat may be slightly

15N-enriched. Plant <515N values are not
influenced by physiogamy or photosyn-
thetic pathway, with one exception. All

Amboseli succulents are 15N-enriched (1 3.2

± 0.5 %c). These species are only distantly
related to each other, and the cause of 15N
enrichment is unclear.

Isotopic Variation in Amboseli Mammals

The C isotope difference between C3
and C4 plants provides a tool for tracing the
diets of Amboseli mammals. Previous field
studies demonstrate a consistent difference
in <513C values between diet and collagen
of ~+5 %o (Vogel 1978; van der Merwe,
1989). Amboseli mammals with pure graz-
ing diets should have <513C values of -8 to
-9 %o. All Amboseli grazers (buffalo, spring
hare, warthog, wildebeest, zebra) have col-
lagen <513C values in this range (Table 26,
Fig. 92B).

In contrast, Amboseli animals with pure
browsing diets should have collagen <513C
values of -22 to -23 %o. None of the
browsers (rninoceros, giraffe) have values
this low, indicating a small but persistent
fraction of grasses or succulents in their

GEOPHYSICAL LABORATORY

167

Table 26. Isotopic data for Amboseli mammals, excluding elephants.

Secies

Common Name

Year of
death

<515N

<5l3C

Carnivores. Insectivores. and Omn

ivores

Crocuta crocuta

Spotted hyena

*1984

16.7

-9.2

Panthera leo

Lion

n.d.

17.1

-10.1

Panthera leo

Lion

n.d.

15.6

-6.4

A cinonyx jubatus

Cheetah

*1984

17.2

-15.7

Canis adustus

Jackal

*1989

15.4

-11.2

Canis adustus

Jackal

*1988

13.8

-9.9

Canis adustus

Jackal

*1988

14.9

-14.8

Otocyon megalotis

Bat-eared fox

*1971

12.2

-10.6

Otocyon megalotis

Bat-eared fox

n.d.

14.6

-12.0

Otocyon megalotis

Bat-eared fox

*1974

14.2

-15.7

Ichneumia albicauda

White-tailed mongoose

n.d.

17.4

-8.3

Orycteropus afer

Aardvark

n.d.

9.7

-12.1

Papio cynocephalus

Yellow baboon

Grazers

1989

10.8

-14.3

Pedetes capensis

Spring hare

*1975

9.8

-9.0

Equus burchelli

Burchell's zebra

1974

9.8

-8.6

Equus burchelli

Burchell's zebra

*1984

9.1

-8.6

Equus burchelli

Burchell's zebra

1990

10.0

-8.5

Connochaetes taurinus

White-bearded wildebeest

1975

12.4

-8.0

Connochaetes taurinus

White-bearded wildebeest

1974

11.0

-8.7

Connochaetes taurinus

White-bearded wildebeest

*1988

13.8

-7.8

Connochaetes taurinus

White-bearded wildebeest

*1989

11.8

-9.0

Syncerus coffer

Buffalo

1968

10.1

-7.9

Syncerus caffer

Buffalo

*1984

10.8

-8.0

Phacochoerus aethiopicus

Warthog

1990

10.8

-8.8

Phacochoerus aethiopicus

Warthog

Browsers

*1989

11.0

-9.2

Dicer os bicornis

Black rhinoceros

1961

6.4

-19.8

Dicer os bicornis

Black rhinoceros

1974

8.2

-18.7

Diceros bicornis

Black rhinoceros

*1984

7.9

-19.0

Giraffa camelopardalis

Giraffe

*1989

11.5

-20.3

Giraffa camelopardalis

Giraffe

Mixed Feeders

*1986

11.6

-19.7

Hystrix cristata

Porcupine

*1973

9.9

-15.6

Hippopotamus amphibius

Hippopotamus

*1973

8.9

-10.0

Hippopotamus amphibius

Hippopotamus

1990

13.1

-8.6

Hippopotamus amphibius

Hippopotamus

*1989

8.9

-10.9

Aepyceros melampus

Impala

1985

12.4

-14.0

Aepyceros melampus

Impala

*1986

11.7

-13.6

Aepyceros melampus

Impala

*1988

13.1

-15.8

Gazella granti

Grant's gazelle

*1988

10.5

-16.0

Gazella granti

Grant's gazelle

*1989

10.2

-15.6

Gazella thomsoni

Thomson's gazelle

1990

14.2

-10.7

Gazella thomsoni

Thomson's gazelle

*1986

10.8

-17.6

* determined by weathering stage

168

w 15

■g

■o

10 11 12 13 14 15 16 17 18

815N

10 -

□ elephant

□ mixed feeder

□ browser
■ grazer
H carnivore

J=L

3 4 5 6 7

10 11 12 13 14 15 16 17 18

CARNEGIE INSTITUTION

Table 27. Isotopic data for Amboseli elephants.

Secimen

Year of
death

<515N «513C

African Elephant: Loxodonta africana

C75-3

1974

12.0

-18.3

C75-6

*1973

9.4

-17.9

E-l )

late 70s

10.4

-13.3

E-3 ]

late 70s

10.3

-13.9

E-8 ]

late 70s

9.7

-17.3

E-ll ]

late 70s

10.7

-13.2

E-13 ]

late 70s

11.4

-13.9

E-15 ]

late 70s

11.9

-13.7

E-17 ]

late 70s

10.3

-15.3

E-19 ]

late 70s

10.1

-13.4

E-21 ]

late 70s

10.4

-14.9

E-23 ]

late 70s

10.4

-11.9

E-25 ]

late 70s

10.2

-15.5

E-27 ]

late 70s

10.4

-15.5

E-29 ]

late 70s

10.3

-13.7

E-30 ]

late 70s

10.7

-11.9

E-34 ]

late 70s

9.8

-14.2

All specimens with Year of death of late 70s were
collected by Cynthia Moss and have known dates
of death that we have not yet received. For Fig.
94 A, these animals are plotted as deaths in 1978.

Fig. 93. (A). Histogram of nitrogen isotope com-
positions for Amboseli plants subdivided accord-
ing to physiogamy. (B) Histogram of nitrogen
isotope compositions for Amboseli mammals sub-
divided according to feeding type. Note that the
difference between the mean £15N values of plants
and animal collagen is only ~+l %o.

diets. Animals known to eat a mixture of
plants (elephant, Grant's and Thomson's
gazelle, hippopotamus, impala, and porcu-
pine) have <513C values intermediate be-
tween browsers and grazers (Tables 26 and
27).

The link between the <513C of diet and
collagen is more difficult to unravel in
carnivores and omnivores, because these
animals obtain carbon both from different
tissues within a body (fat, muscle, skin) and

from plants. All these sources may have
different <513C values. Generally, herbi-
vore meat and carnivore collagen differ by
~ +5%o, but the difference between herbi-
vore collagen and carnivore collagen is
+2%c (van der Merwe, 1989). Observa-
tions of hunting carnivores suggest that
hyena and lion consume chiefly C4-feed-
ing herbivores (wildebeest and zebra),
whereas cheetah eat mixed feeders
(Thomson's and Grant's gazelle and im-
pala). These observations are supported by
#3C values (Table 26, Fig. 92B). The
smaller carnivores (fox, jackal, mongoose)
eat smaller animals from across the dietary
spectrum and exhibit a spread of <513C
values.

GEOPHYSICAL LABORATORY

169

IB -
16-

a Carnivores

• Browser & Mixed

14-

10
to

12-
10-

Carnivore A •

•

• *"

Browser & Mixed

•
•

1 1 1 r

1955 1960 1965 1970 1975 1980 1985 1990 1995

Year of Death

16
14
12
10 H

£ 8
to

4-
2-

-f-

hT~'teHrH

-30

-25

I I '» I | ill l

-20
813C

-15

-10

A Rhinoceros
• Elephant
*±* Mean & std. dev. of
African park elephants

Fig. 94. (A) Secular variation in the nitrogen
isotope composition of mammals. Grazers: Y=
-68.25 +.0.04X r=0.24, slope is not significantly
different from 0. Data and regression line are not
plotted. Browsers: Y= -279.33 + 0.15X r=0.62,
slope is significantly different than 0. Carnivores:
Y= -253.66 + 0. 14X r=0.58, slope is significantly
different than 0. The elephants listed as late 70s
deaths are included on the figure, and given 1978
as the year of death, but they were not used in the
regression calculation. (B) Carbon and nitrogen
isotopic composition of Amboseli elephants and
rhinoceros. Also plotted are the means and stan-
dard deviations for elephant ivory from 16 other
African parks and preserves. Firm determination
of temporal isotopic trends within species must
await analysis of specimens with a greater range of
known ages of death. However, the spread in the
isotopic data from both species may result from
coupled increases in <513C and <515N values with
time. Although values from Amboseli elephants
do not overlap with values from many other parks,
they vary by amounts as great as those used to
discriminate between other park populations.

A fractionation of ~ +3%c between the
<515N value of plant food and herbivore
collagen has been reported in previous
studies (DeNiro and Epstein, 1981 ; Hare et
ai, 1991). Given this fractionation, the
average Amboseli herbivore should have a
collagen Sl 5N value of 1 2- 1 3 %o. Although
some animals have values in this range,
most are more negative (Tables 26 and 27,
Fig. 93B) Indeed, using the mean ^5N of
animals and a fractionation of +3 %o, we
would expect dietary plants with values of
l%o or less. Few plants within the park
have isotopic values this low.

There are several plausible explana-
tions for this discrepancy. First, we ana-

lyzed only plants collected in a dry season.
In the wet season, plants may have lower
<515N values. Second, if the 515N value of
plants from Amboseli has increased re-
cently, current vegetation may not be repre-
sentative of the foods eaten by the sampled
animals . Finally, the fractionation between
diet and herbivore collagen may not be
+3%o. Variability in this fractionation has
been detected previously (Ambrose and
DeNiro, 1986; Heatonef a/., 1986; Sealy et
al., 1987), and attributed to differences in
N metabolism between different herbivores.
A fractionation of ~+\%o is observed be-
tween current Amboseli plants and the
sampled herbivores.

170

CARNEGIE INSTITUTION

Differences in collagen <515N between
herbivores and carnivores are well studied
and range from +3 to +6 %o (Schoeninger
and DeNiro, 1984; Ambrose and DeNiro,
1986; Sealy et al., 1987). Amboseli graz-
ers, mixed feeders and browsers averaged
10.9 ± 1.3 %o, 9.1 ± 2.3 %o9 and 10.9 ± 1.3
%o, respectively, whereas true Amboseli
carnivores averaged 15.4 ± \.6%c. Conse-
quently, there is a trophic level fraction-
ation of ~+5 %o. The omnivorous yellow
baboon has a lower <515N value, suggesting
a preponderance of plant foods in the diet.
Finally, the aardvark, which consumes ants
and termites, has a <515N value within the
herbivore range. However, insect chitin is
known to be 15N-depleted relative to di-
etary plants (Schimmelmann,pers. comm.),
which would lead to lower values in the
collagen of insectivores relative to carni-
vores.

Secular Trends in the Isotopic Composi-
tion of Amboseli Mammals

The Amboseli ecosystem has changed
dramatically since 1960 because of a loss
of trees and the expansion of grassland. To
document isotopic trends in park mammals,
multiple samples of individual species from
different time periods must be examined.
Our data are currently insufficient for such
a treatment, but trends within broad feed-
ing categories may be examined. Collagen
<515N and <513C values, and thus the diet of
grazers, have remained constant from 1968
to 1990 (Table 26). The &$N of browsers
and mixed feeders has increased by a sta-
tistically significant amount (Fig. 94A).

This trend, however, is strongly influenced
by a low value for a single old specimen.
Although most elephant deaths are only
roughly dated to the late 1 970s, the popula-
tion seems to be trending towards higher
#5n and 8&C values (Fig. 94B). The
magnitude of this variation is significant
when compared to the differences between
populations from different parks. There is
a suggestion of similar coupled increases
for rhinoceros, but the sample is quite small.
Finally, carnivores also increase in <515N
with time by amount similar to browsers
(Fig. 94A).

We hypothesize that as the park was
stripped of trees, browsers and mixed feed-
ers have been forced to consume more
grass. Increased grass consumption is par-
ticularly evident for elephants. However,
grass is relatively nutrient-poor compared
to browse. Eventually, the browsers and
mixed feeders suffered nutritional stress.
Nutritional stress may cause an animal to
remetabolize previously deposited proteins,
and can potentially produce an increase in
collagen <515N in bones equivalent to that
generated by feeding at a higher trophic
level (Tuross, pers. com.). The grazers
thrived as the grasslands expanded and
exhibit no isotopic changes. Carnivores
eat both types of herbivores, and conse-
quently exhibit intermediate isotopic trends.

Conclusions

The carbon and nitrogen isotopes in
most plants from Amboseli National Park
varied as expected, with a strong differen-
tiation in 5J 3C between C3, C4, and aquatic

GEOPHYSICAL LABORATORY

171

plants, and a high mean d^N value. The
15N enrichment of succulent plants was
unexpected, and is currently unexplained.
Differences in the <513C of plants is re-
flected in the collagen of the animals that
consume these plants, and ultimately can
be detected at higher trophic levels when
these herbivores are preyed upon by carni-
vores. The fractionations of C and N iso-
topes between diet and collagen that we
discovered match previous reports, with
one exception. The fractionation of N
between plant and herbivore collagen was
much lower than expected. Finally, al-
though our observations must be supported
by more extensive sampling, there are sta-
tistically significant secular trends in the
<515N of Amboseli browsers and mixed
feeders and carnivores, whereas grazers
are invariant. The substantial isotopic trends
shown by Amboseli elephants may indi-
cate that stable isotopes will be of limited
utility in tracing the source of elephant
ivory in changing habitats.

References

Ambrose, S. H., and M. J. DeNiro, The isotopic
ecology of East African mammals, Oecologia,
69, 395-406, 1986.

Behrensmeyer, A. K., Taphonomic and ecologic
information from bone weathering, Paleobio.,
4, 150-162, 1978.

DeNiro, M. J., and S. Epstein, Influence of diet on
the distribution of carbon isotopes in animals,
Geochim. Cosmochim. Acta, 42, 495-506,
1978.

DeNiro, M. J., and S. Epstein, Influence of diet on
the distribution of nitrogen isotopes in ani-
mals, Geochim. Cosmochim. Acta, 45, 341-
351, 1981.

Hare, P. E., M. L. Fogel , T. W. Stafford, Jr., A. D.
Mitchell, and T. C. Hoering, The isotopic
composition of carbon and nitrogen in indi-
vidual amino acids isolated from modern and

fossil proteins, J. Archaeol. Sci, in press, 199 1 .

Heaton, T. E., J. C. Vogel, G. von la Chevallerie,
and G. Collett, Climatic influence on the iso-
topic composition of bone nitrogen, Nature,
322, 822-823, 1986.

Raven, J. ., The application of mass spectrometry
to biochemical and physiological studies, in
The Biochemistry of Plants, Vol. 13, Aca-
demic Press, Inc., New York, pp. 127-180,
1987.

Schoeninger, M. J., and M. J. DeNiro, Nitrogen
and carbon isotopic composition of bone col-
lagen from marine and terrestrial animals,
Geochim. Cosmochim. Acta, 48, 625-639
1984.

Sealy, J.C., N. J. van der Merwe, J. A. Lee Thorp,
and J. L. Lanham, Nitrogen isotopic ecology
in southern Africa: implications for environ-
mental and dietary tracing, Geochim.
Cosmochim. Acta, 57,2707-2717, 1987.

Tieszen, L. L., and T. W. Boutton, Stable carbon
isotopes in terrestrial ecosystem research, in
Stable Isotopes in Ecological Research,
Rundel, P.W., J.R. Ehleringer, and K. A. Nagy ,
eds., Springer- Verlag, New York, pp. 167-
195, 1988

Tuross, N.,M. L. Fogel, andP. E. Hare, Variability
in the preservation of the isotopic composition
of collagen from fossil bone, Geochim.
Cosmochim. Acta, 52, 929-935, 1988.

van der Merwe, N. J., Natural variations in 13C
concentration and its effect on environmental
reconstruction using 13C/12C ratios in animal
bones, in The Chemistry of Prehistoric Human
Bone, Price, T.D., ed., Cambridge Univ. Press,
New York, pp. 105-125, 1989.

van der Merwe, N. J., J. A. Lee-Thorp, J. F.
Thackeray, A. Hall-Martin, F. J. Kruger,
H.Coetzee, R. H. V. Bell, and M. Lindeque,
Source-area determination of elephant ivory
by isotopic analysis, Nature, 346, 744-746,
1990.

Vogel, J. C, Isotopic assessment of the dietary
habits of ungulates, S. Afr. J. Sci., 74, 298-30 1 ,
1978.

Vogel, J. C, B. Eglington, and J. M. Auret, Isotope
fingerprints in elephant bone and ivory, Na-
ture, 346, 747-749, 1990.

172

CARNEGIE INSTITUTION

Rapid Racemization of Aspartic Acid

in mollusk and ostrich eggshells :

A New Method for Dating on

a Decadal Time Scale

Glenn A. Goodfriend, David W. von
Endt* and RE. Hare

Epimerization of L-isoleucine to D-
alloisoleucine has been extensively ana-
lyzed in mollusk shells as a means of deter-
mining relative or absolute ages, primarily
of Pleistocene samples. More recently,
epimerization in Pleistocene ostrich egg-
shells has been studied (Brooks et al. , 1 990).
Racemization of a number of other amino
acids has been studied in mollusk shells,
leucine being the most widely studied (e.g.,
Wehmiller, 1984; reviewed in Goodfriend,
1991). Until recently, research on aspartic
acid racemization in mollusks had been
limited to some Pleistocene marine samples
from the West Coast of the U. S.
(Kvenvolden et al., 1979). But recent stud-
ies on aspartic acid racemization in desert
land snails have turned up an interesting
pattern: the initial rate of racemization is
extremely high; the rate then slows down
progressively with increasing time or D/L
ratio. This pattern has been demonstrated
in both a radiocarbon-dated series of Holo-
cene shells (Goodfriend, 1 99 1 ) as well as in
heating experiments of modern shells
(Goodfriend and Meyer, 1991).

The very rapid initial rate of racemiza-
tion presents the possibility of using aspar-

Conservation Analytical Laboratory,
Smithsonian Institution, Washington, D.C. 20560

tic acid racemization as a high-resolution
dating method for young materials. This is
of particular interest because radiocarbon
generally cannot be used for dating of post-
1650 A.D. samples (see radiocarbon cali-
bration curve of Stuiver and Pearson ( 1 986)
for this period). On the other hand, evalua-
tion of the precision of aspartic acid racem-
ization dating in this time range is difficult,
because of the unavailability of radiocar-
bon as an independent measure of age. For
this reason, we turned to museum mollusk
collections as a source of material of known
age with which to evaluate the age predic-
tive ability of D/L aspartic acid ratios. In
addition, we examined aspartic acid race-
mization in three samples of ostrich egg-
shells, to see if the phenomenon of rapid
initial racemization also occurs in this ma-
terial.

Materials and Methods

Seven samples of the land snail
Triodopsis multilineata from Iowa and
Kansas were obtained from the collection
of the U. S. National Museum of Natural
History. The dates of collection of these
samples range from 1881 to 1949. Most of
these could be seen to have been collected
alive because of the presence of some re-
mains of the animals inside the shells. In
addition, two modem samples (1990 and
1991) of another Triodopsis species, T.
tridentata, were collected in the Chevy
Chase district of Washington, D. C. Samples
of modem ostrich eggs hatched at the Front
Royal, Virginia, breeding farm of the Na-
tional Zoological Park in 1978 and at Dolly

GEOPHYSICAL LABORATORY

173

Farms (Vicksburg, Mississippi) in 1990
were obtained. A bead fashioned out of
ostrich eggshell, excavated from a fort at
Oudepost, South Africa (occupied ca. 1652-
1660 A.D.; Schrire et al., 1990) was also
obtained for analysis. (The authors are in-
debted to R. Herschler and P. Greenhall of
the U. S. National Museum of Natural
History for providing the material of
Triodopsis multilineata used in this study.
Ostrich egg samples were kindly provided
by A. Brooks, C. Schrire, J. Kokis, the
National Zoological Park, and R. Shafer of
Dolly Farms.)

The periostracum (the outer organic
layer, covering or partly covering the
mollusk shells) was ground off the shell
samples by abrasive-tipped bits using a
motorized hand-held tool. Organic mate-
rial on the surface of the ostrich egg pieces
was similarly removed. Samples were then
subjected to a short dip in dilute HC1,
washed three times in distilled water, and
dried under vacuum. Hydrolysis and
derivatization of the amino acids to their N-
trifluoracetyl isopropyl ester derivatives
were carried out as described in Goodfriend
(1991). The D/L amino acid ratios were
analyzed by gas chromatography using a
Hewlett-Packard model 5790. The values
are reported as the ratio of the areas of the
D and L peaks, as calculated by a Hewlett-
Packard model 3 3 94 A integrator. For the
Triodopsis samples, analyses were based
on preparations made from small pieces of
three individual shells (total weight: 13-19
mg); a piece from a single shell was ana-
lyzed in the case of the T. tridentata and
ostrich egg samples. At least two analyses
of each snail shell preparation were carried

out and the mean of these is reported. The
standard error of these sample means aver-
aged 0.001 1. The ostrich eggshell samples
were analyzed once.

Results

The Triodopsis shells show a progres-
sive increase in the D/L aspartic acid ratio
with increasing age (Fig. 95). The initial
value of 0.041 to 0.044 represents either
racemization induced by the preparation
procedure or the occurrence of small
amounts of D-aspartic acid in the modern
shells. From this initial value, the D/L ratio
increases up to 0.093 in the 110-year-old
specimens (collected in 1881). Although
the modem specimens are of a different
species than the others, different mollusk
species within the same genus generally
show the same rates of epimerization (Lajoie
et al., 1980), so it is expected that the
modem T. tridentata values are represen-
tative of those of modem T multilineata. A
simple linear regression of the D/L ratios
on the age yields the equation

D/L = 0.000458 (age) + 0.00431,

which indicates an average racemization
rate of l%/22 yr. (In this and subsequent
analyses, a single anomalous sample (D/L
of 0. 12 for a 1949 sample) was left out; this
high value may have been the result of
heating of the shell to extract the bodies or
to dry the shells after extraction of the
bodies.) The correlation coefficient between
the D/L ratio and age is 0.979. An estimate

174

CARNEGIE INSTITUTION

0.10

-i — i — i — i — i-

■ T. multilineata
• T. tridentata

1990 1970 1950 1930 1910 1890

year A. D.

Fig.95. DA- aspartic acid ratios in shell samples of
two species of land snails of the genus Triodopsis.
The year of collection of the samples is indicated
on the horizonal axis.

of the error of an age predicted from the D/
L ratio of a specimen was calculated from
the data as the square root of the mean
square error of a regression of age on D/L
ratio. A value of 9.5 yr was thus obtained
and indicates that the year of collection of
an undated shell sample can be estimated
with approximately this degree of precision
based on analysis of its D/L aspartic acid
ratio. This estimate assumes that the scatter
of the points about the regression line is
uniform over the range of values analyzed,
whereas it appears that the scatter is greater
at higher ratios, which is as expected if the
scatter is due to variation in the average
racemization rate of different samples. Thus
the error of age estimations would be lower
than the calculated value for more recent
ages and higher for older ages.

The few results available for the ostrich
egg samples (Table 28) also suggest a high
aspartic acid racemization rate in this ma-

terial. A net racemization of about 0.095 , or
a rate of about 1% racemization per 35
years, is indicated for the eggshell sample
about 330 years old.

Discussion

The results confirm the occurrence of a
very high rate of aspartic acid racemization
in museum mollusk material, as expected
based on earlier studies of desert land snails.
Because the samples show a regular pattern
of increasing D/L ratios with increasing
age, aspartic acid racemization may be
useful as a dating method for materials on
a decadal time scale. For study of museum
collections, this may have several applica-
tions. In biogeographical studies, it is often
of importance to know when a specimen or
set of specimens was found at a particular
location, since distributions may change
over short time scales. Aspartic acid race-
mization analysis could be used to deter-
mine the approximate time of collection of
undated samples in museum collections. It
could also be used to determine if speci-
mens were alive (or freshly dead) when
collected, or whether they represent older,
dead-collected material. Older records of
distributions could be obtained from such
material. The method may also be applied

Table 28. D/L aspartic acid ratios in some ostrich
eggshell samples.

Year

Source

D/L

1990 Dolly Farms 0.044

1978 Natl. Zoological Park 0.057

ca. 1652-1660 Oudepost, S. Afr. 0.129

GEOPHYSICAL LABORATORY

175

to dating of recent deposits in nature and
should provide good time resolution for the
post- 1650 A.D. period not covered by ra-
diocarbon. Such applications require an in
situ rate calibration based on independentiy-
dated material of the same species. An in
situ calibration is required since different
museum collections or different field sites
will differ in their average temperatures. In
museum collections, this calibration can be
obtained from other specimens of known
collection dates. In the field, radiocarbon
dates from pre- 1650 A.D. samples or dat-
ing by association with archeological arti-
facts are the most likely sources of cali-
bration dates. Possible problems with ap-
plication of the method to museum mate-
rials may arise if samples have been sub-
jected to prolonged heating or high tem-
peratures during processing or storage.

Aspartic acid racemization has been
applied previously to dating of human teeth
(e.g., Helfman and Bada, 1976; Ohtani et
al., 1988), where it shows a high rate of
racemization (1% per 13 yr in dentine;
Helfman and Bada, 1976). However, this
high rate occurs at body temperature, or
about 37°C. The projected rate at room
temperature (about 2 1 °C) would be 14 times
slower, or 1% per 180 yr (assuming an
activation energy of 30 kcal/mol). Thus
very rapid racemization of aspartic acid in
young materials seems to be limited to
biogenic carbonate materials, such as mol-
lusk shells and bird eggshells; biogenic
phosphates, as represented by teeth, show a
considerably slower rate. One may expect
the phenomenon of very rapid initial race-
mization of aspartic acid to be found also in
other biogenic carbonates, such as fora-

miniferal tests and coral skeletons.

Some possible modifications of ana-
lytical procedures could lead to even higher
temporal resolution. For example, it has
been found that the free amino acid fraction
in mollusks is always more highly
epimerized than the total amino acid frac-
tion that is obtained from hydrolysis (e.g.,
Miller and Hare, 1980). Although D/L en-
antiomer ratios have never been measured
in the free amino acid fraction, it might be
expected that this would yield similar re-
sults, i.e., that the DA- aspartic acid ratio
may increase faster in the free amino acid
fraction than in the total. It has been sug-
gested that the rapid initial rate of aspartic
acid racemization may actually be the re-
sult of the racemization of asparagine rather
than aspartic acid per se (Goodfriend, 1 99 1 );
asparagine is converted to aspartic acid
during hydrolysis, so what is measured as
"aspartic acid" is actually the sum of the
aspartic acid and asparagine originally
present in the sample. Development of
methods for measuring the D/L ratio of
asparagine may result in better time resolu-
tion for dating applications.

References

Brooks, A. S., P. E. Hare, J. E. Kokis, G. H. Miller,
R. D. Ernst, and F. Wendorf, Dating Pleisto-
cene archeological sites by protein diagenesis
in ostrich eggshell, Science, 248, 60-64, 1990.

Goodfriend, G. A., Patterns of racemization and
epimerization of amino acids in land snail
shells over the course of the Holocene,
Geochim. Cosmochim. Acta, 55, 293-302,
1991.

Goodfriend, G. A., and V. R. Meyer, A compara-
tive study of amino acid racemization/
epimerization kinetics in fossil and modern
mollusk shells, Geochim. Cosmochim. Acta,
in press.

176

CARNEGIE INSTITUTION

Helfman, P. M., and J. L. Bada, Aspartic acid
racemisation in dentine as a measure of age-
ing, Nature, 262, 279-281, 1976.

Kvenvolden, K. A., D. . Blunt, andH. E. Clifton,
Amino-acid racemization in Quaternary shell
deposits at Willapa Bay, Washington,
Geochim. Cosmochim. Acta, 43, 1505-1520,
1979.

Lajoie, K. R., J. F. Wehmiller, and G. L. Kennedy,
Inter- and intrageneric trends in the apparent
racemization kinetics of amino acids in Qua-
ternary mollusks, mBiogeochemistry of Amino
Acids, P. E. Hare, T. C Hoering, and K. King,
Jr., eds., John Wiley and Sons, New York, pp.
305-340, 1980.

Miller, G. H. and P. E. Hare, Amino acid geochro-
nology: integrity of the carbonate matrix and
potential of molluscan fossils, mBiogeochem-
istry of Amino Acids, P. E. Hare, T. C. Hoering,
and K. King, Jr., eds., John Wiley and Sons,
New York, pp. 415-443, 1980.

Ohtani, S., S. Kato, H. Sugeno, H. Sugimoto, T.
Marumo, M. Yamazaki, and K. Yamamoto, A
study on the use of the amino-acid racemiza-
tion method to estimate the ages of unidenti-
fied cadavers from their teeth, Bull. Kanagawa
Dental College, 16, 11-21, 1988.

Schrire, C, J. Deetz, D. Lubinsky, and C.
Poggenpoel, The chronology of Oudepost I,
Cape, as inferred from an analysis of clay
pipes, J. Archaeol. ScL, 17, 269-300, 1990.

Stuiver, M. and G. W. Pearson, High-precision
calibration of the radiocarbon time scale, AD
1950-500 BC, Radiocarbon, 28, 805-838,
1986.

Wehmiller, J. F., Relative and absolute dating of
Quaternary mollusks with amino acid racem-
ization: evaluation, applications and questions,
in Quaternary Dating Methods, W. C.
Mahaney, ed., Elsevier, Amsterdam, pp. 171-
193, 1984.

A Burning Question: Differences

between Laboratory-Induced

and Natural Di agenesis in Ostrich

Eggshell Proteins*

A. S. Brooks, RE. Hare, J.E. Kokis, and
K. Durana

In earlier papers (Brooks et ai, 1990;
Kokis et al., 1990), we demonstrated the
utility of the D-alloisoleucine/L-isoleucine
(A/I) ratio in ostrich eggshell for estimat-
ing the age of archaeological specimens.
Of the biogenic carbonates and phosphates
tested so far, ostrich eggshell most nearly
approximates a closed system with little
loss of either water or protein breakdown
products to the environment. Although
ostrich eggshell is a common material in
archaeological sites located in the arid and
semiarid regions of the Old World, human
activity and natural factors at these sites
may produce anomalous epimerization ra-
tios in two ways: by heating (camp fires,
brush fires) and by stratigraphic mixing
through ancient excavations (pits, burials,
burrows, etc.) into underlying deposits
which may also contain eggshell from hu-
man occupation debris. For any archaeo-
logical horizon, it is important to be able to
distinguish between heating and strati-
graphic admixture, especially of older ma-
terials. Heating to certain higher tempera-
tures may indicate the presence of human-
controlled fire, whereas stratigraphic ad-
mixture may call into question the interpre-

This work is supported by NSF Grant BNS-
9011657

GEOPHYSICAL LABORATORY

177

tation of other materials at the site, e.g.,
human fossils. In addition, if temperature
differentially affects two decomposition
reactions because they have different acti-
vation energies, we may be able to use
these two reactions to determine simulta-
neously both time and temperature. In this
paper, we describe results from laboratory
heating of ostrich eggshell fragments at
controlled temperatures for varying peri-
ods. Amino acid compositions of these
heated samples were compared to amino
acid compositions of archaeological
samples from the last 80,000 years.

Materials and Experimental Methods

Heating experiments were conducted
in a heated aluminum block with a Model
71 A Temperature Controller (RFL Indus-
tries, Inc., Boonton, New Jersey). Tem-
perature readings were within ±0.2° C.
Samples of eggshell fragments were
weighed and dropped into pre -heated tubes
placed in the aluminum block. Samples
were then processed for free and total amino
acids as described by Brooks et al. (1990).

Three different sample series of ostrich
eggshell fragments were heated in the labo-
ratory, and two sets of archeological samples
were analyzed. (Modem ostrich eggshell
samples were provided by the National
Zoological Park, Washington, D.C. and R.
Shafer of Dolly Farms, Vicksburg, Missis-
sippi. Archaeological eggshell samples
were provided by O. Bar-Yosef, A.S.
Brooks, J. Deacon, M. Mehlman, and W.E.
Wendt.)

Laboratory-Heated Samples:

(1) A series heated dry at 300°C for
incremental time periods of 15 min., 30
min., 1 hr.,2hrs.,4hrs., 8hrs., 16hrs.,and
32hrs.

(2) A series heated dry for one hour
each at temperatures of 160°C, 200°C,
240°C, 280°C, 320°C, and 360°C.

(3) A series heated for incremental
time periods in water vapor at 157°C at
times ranging from 0.5 to 256 hours.
Archaeological Samples:

(4) A stratified series of archaeological
samples from a tropical zone site (7uGi,
Botswana), in which the A/I ratio increases
regularly with age and depth to above 1.0.

(5) A group of pieces with anomalous
A/I ratios from archaeological sites (7iGi,
Boomplaas, Mumba Shelter, Qafzeh,
Apollo 11). These pieces either had no
significant A/I peaks or had ratios which
were much higher than others from the
same or underlying levels.

For each series we measured the ratios
of A/I peak areas subtracting 0.015 from
each ratio to correct for laboratory-induced
epimerization. We also measured peak ar-
eas of aspartic acid (Asp), glutamic acid
(Glu), glycine (Gly), alanine (Ala), and
ammonia (NH3), as well as serine (Ser),
threonine (Thr), and arginine (Arg).

Results

In the heating experiments, a sequence
of changes in amino acid composition and
concentrations occurs that can be described
as a series of stages.

178

CARNEGIE INSTITUTION

Stage 0: Modern unheated. Four major
amino acids (Glu, Gly, Asp, Ala) occur in
comparable amounts. Other amino acids,
including Ser, Arg, Thr, and He, are also
present at lower levels. No significant
amounts of alloisoleucine are found. The
level of NH3 relative to the four major
amino acids is insignificant.

Stage 1. Light heating (1 hour at 160°-
200° C). The four major amino acids persist
in comparable amounts, whereas Ser, Arg,
and Thr diminish to trace levels.
Alloisoleucine increases with length of heat-
ing. The NH3 level is elevated.

Stage 2. Moderate heating (1 hour at
200°C-280°C). The four major amino ac-
ids no longer remain at comparable levels;
glutamic remains relatively constant while
the others decrease. Serine, threonine, and
arginine diminish to only trace levels or are
completely absent. A/I values are not al-
ways possible to determine due to the ap-
pearance of interfering peaks. The level of
NH3 steadily increases from 200°C to 280°C
and over time.

Stage 3. Strong heating (1 hour at
300°C-360°C). The four major amino ac-
ids diminish to only trace levels. Some Glu
persists after other amino acids disappear.
NH3 is the predominant peak. Some new
peaks appear at this stage, which are tenta-
tively identified as y-amino-butyric acid
(GABA), and some amines, possibly
methyl, ethyl, and propyl. Alloisoleucine
and isoleucine levels are too low to calcu-
late with confidence. Interestingly, there
appears to be some synthesis of amino
acids at these low levels.

In the archaeological series the same
four stages are observed. Samples that show
good correlations with other age estimates,
such as radiocarbon dating, exhibit only
Stage 0 or 1 patterns. Some of the anoma-
lous samples from archaeological sites that
do not correlate with the other age esti-
mates show stage 2 or stage 3 patterns, with
high NH3 levels and the presence of amines
and probably GABA. Other anomalous ar-
chaeological samples exhibit little change
from stage 0 or early stage 1 patterns; these
show no evidence of heating and are pre-
sumed to have derived from underlying
levels by stratigraphic admixture.

Discussion

Differences were immediately apparent
on inspection of the chromatograms for the
archaeological series compared to the
heated series. In the archaeological series,
little significant decrease was noted in the
four "stable" amino acids studied, even
while A/I ratios increased to over 1 .0. In
addition, there was little build-up of NH3.
In the heated samples, in contrast, at tem-
peratures as low as 200°C for 1 hour, or at
157°C for 32 hours, there is a net increase
in NH3 and a decrease in aspartic concen-
tration relative to the more stable concen-
tration of glutamic acid. The archaeologi-
cal samples from 7cGi differ from all the
pieces from series 2 heated at 200°C or
above, all the pieces from series 1 (300°C),
and all pieces from series 3 (157°C) heated
32 hours or more in two respects: (1) heated

GEOPHYSICAL LABORATORY

179

pieces exhibit a higher concentration of
NH3 than of aspartic acid, and (2) no ar-
chaeological piece has more than 50% as
much NH3 as aspartic acid, except for the
anomalous pieces.

Of the anomalous archaeological pieces
studied, three high A/I ratio pieces from
7iGi were found in the top Later Stone Age
horizons where A/I ratios normally ranged
from 0.1 to 0.5 and radiocarbon calibra-
tions suggested an age in the last 35,000
years. Two of these pieces clearly fit with
the earlier Middle Stone Age series and are
presumably derived from below by human
or animal disturbance. However, one of the
three which provided an infinite radiocar-
bon age (>40,000 yr B.R) must have been
heated in antiquity, as it matched the amino
acid composition of the strongly heated
(early Stage 3) laboratory samples. Two
anomalous pieces from the top levels at
another site (Boomplaas) gave A/I ratios
higher than the pieces from the bottom
levels whose estimated age was 80,000
B.R One of these pieces has been dated by
TAMS 14C to 5220 ± 70 yr B.R It exhibits
an NH3-to-aspartic acid ratio greater than
1.0, and is presumed to have been heated.

Other anomalous archaeological pieces
contained high NH3 concentrations but
insufficient amounts of alloisoleucine or

isoleucine to establish the A/I ratio. As
mentioned above, experimental pieces sub-
jected to stronger heating conditions (320°C
for one hour, or 300° C for four or more
hours) exhibited high NH3 and g-amino-
butyric acid peaks as well as other very
small peaks. A/I ratios in these strongly
heated pieces could not be measured pre-
cisely, but did not appear to progress much
above 0.7. Examination of the small peaks
suggests that at the higher temperatures in
the heating experiments, serine as well as
some other amino acids are being synthe-
sized. Archaeological samples from Apollo
11 in Namibia, Mumba Shelter in Tanza-
nia, and Qafzeh Cave in Israel also con-
form to this latter pattern and were almost
certainly heated to relatively high tempera-
tures in antiquity.

References

Brooks, A.S.,P.E. Hare, J. E. Kokis, G. H. Miller,
R. D. Ernst, andF. Wendorf, Dating pleistocene
archaeological sites by protein diagenesis in
Ostrich eggshell, Science, 248, 60-64, 1990.

Kokis, J. E., A. S. Brooks, andP. E. Hare, Chronol-
ogy a nd aminostratigraphy of Middle and Late
Stone Age sites from Sub-saharan Africa: A
comparison of protein diagenesis and radio-
carbon dating of ostrich eggshell, Geological
Society of America Abstracts With Programs,
22, A145-146, 1990,

GEOPHYSICAL LABORATORY

181

Publications

Reprints of the numbered publications listed below are available, except where noted, at no charge
from the Librarian, Geophysical Laboratory, 5251 Broad Branch Road, N.W, Washington, D.C.
20015-1305, U.S.A. Please give reprint number(s) when ordering. Youmay also request to be placed
on the Laboratory's mailing list to receive periodic notifications of recent publications.

Angel, R. J., N. L. Ross, L. W. Finger, and R. M.
Hazen, Ba3CaCuSi60n: A new {1B,1s(i,oo)}
{4Si60i7} chain silicate, Acta Crystallogr.
C46, 2028-2030, 1990 (G.L. Paper 2190).

Angel, R. J., R. K. McMullen, and C. T. Prewitt,
Substructure and superstructure of mullite by
neutron diffraction, Am. Mineral., 76, 332-
342, 1991 (G.L. Paper 2216).

B adding, J. V., H. K. Mao, and R. J. Hemley,
High-pressure synchrotron X-ray diffraction
of Cs IV and Cs V, Solid State Commun., 77,
801-805, 1991 (G.L. Paper 2208).

Bebout, G. E., Field-based evidence for
devolatilization in subduction zones: Implica-
tions for arc magmatism, Science, 251, 413-
416, 1991 (G.L. Paper 2206).

Bebout, G. E., Geometry and mechanisms of fluid
flow at 15 to 45 kilometer depths in an early
cretaceous accretionary complex, Geophys.
Res. Lett., 18, 923-926, 1991 (G.L. Paper
2217).

Chamberlain, C. P., J. M. Ferry, and D. Rumble,
III, The effect of net-transfer reactions on the
isotopic composition of minerals, Contrib.
Mineral. Petrol., 105, 322-336, 1990 (G.L.
Paper 2192; no reprints available for distribu-
tion).

Cifuentes, L. A., L. E. Schemel, and J. H. Sharp,
Qualitative and numerical analysis of the ef-
fects of river inflow variations on mixing
patterns in estuaries, Estuarine Coastal Shelf
Sci., 30, 41 1-427, 1990 (G.L. Paper 2198; no
reprints available for distribution).

Coffin, R. B., D. J. Velinsky, R. Devereux, W. A.
Price, and L. A. Cifuentes, Stable carbon iso-
tope analysis of nucleic acids to trace sources
of dissolved substrates used by estuarine bac-

teria, Appl. Environ. Microbiol., 56, 2012-
2020, 1990 (G.L. Paper 2191).

Cohen, R. E., Bonding and elasticity of stishovite
Si02 at high pressure: linearized augmented
plane wave calculations, Am. Mineral., 76,
733-742, 1991 (G. L Paper 2220).

Fei, Y., H. K. Mao, and B. O. My sen, Experimen-
tal determination of element partitioning and
calculation of phase relations in the MgO-
FeO-SiC>2 system at high pressure and high
temperature, /. Geophys. Res., 96, B2, 2157-
2169, 1991 (G.L. Paper 2205).

Finger, L. W., R. M. Hazen, and C. T. Prewitt,
Crystal structures of Mgi2SUOi9(OH)2 (Phase
B) and Mgi4Si5<I>24 (Phase AnhB), Amer.
Mineral, 76, 1-7, 1991 (G.L. Paper 2219).

Hare, P. E., M. L. Fogel, T. W. Stafford, Jr., A. D.
Mitchell, and T. C. Hoering, The isotopic
composition of carbon and nitrogen in indi-
vidual amino acids isolated from modern and
fossil proteins, /. Archaeol. Sci., 18, 277-292,
1991 (G.L. Paper 2215).

Hanfland, M., R. J. Hemley, and H. K. Mao,
Optical absorption measurements of hydrogen
at megabar pressures, Phys. Rev. B, 43, 8767-
8770 1991 (G.L. Paper 2213).

Hazen, R. M., Crystal structures of high-tempera-
ture superconductors, in Physical Properties
of High-Temperature Superconductors II, D.
M. Ginsberg, ed., Chapter 3, pp. 121-198,
World Scientific, New Jersey, 1990 (G.L. Pa-
per 2158; no reprints available for distribu-
tion).

Hazen, R. M., and J. S. Trefil, Science Matters:
Achieving Scientific Literacy, Doubleday, New
York, 1991 (G.L. Paper 2195) (Available at
your local bookstore or if you prefer direct
from Doubleday)

182

CARNEGIE INSTITUTION

Hazen, R. M., and J. S. Trefil, Achieving geologi-
cal literacy, /. Geol. Educ, 39, 28-30, 1991
(G.L. Paper 2196)

Hazen, R. M, J. Zhang, and J. Ko, Effects of Fe/
Mg on the compressibility of synthetic
wadsleyite:P-(Mg1.J^ejt)2Si04(x<Q.25),P/ry5.
Chem. Minerals ,17, 416-419, 1990 (G.L. Pa-
per 2197).

Hemley, R. J., H. K. Mao, L. W. Finger, A. P.
Jephcoat, R. M. Hazen, and C. S. Zha, Equa-
tion of state of solid hydrogen and deuterium
from single-crystal X-ray diffraction to 26.5
GPa, Phys. Rev. B, 42, 6458-6470, 1990 (G.L.
Paper 2180).

Hemley, R. J., andH. K. Mao, Critical behavior in
the hydrogen insultator-metal transition, Sci-
ence, 249, 391-393, 1990 (G.L. Paper 2184).

Hemley, R. J., H. K. Mao, and M. Hanfland,
Spectroscopic investigations of the insulator-
metal transition in solid hydrogen, in Molecu-
lar Systems under High Pressure (Proceed-
ings of the II Archimedes Workshop on Mo-
lecular Solids under Pressure Catania, Italy,
28-31 May 1990) R. Pucci, and G. Piccitto,
eds., pp. 223-243, Elsevier, New York, 1991
(G.L. Paper 2189).

Hemley, R. J., H. K. Mao, and J. F. Shu, Low-
frequency vibrational dynamics and structure
of hydrogen at megabar pressures, Phys. Rev.
Lett., 65, 2670-2673, 1990 (G.L. Paper 2201).

Hemley, R. J., and J. D. Kubicki, Deep mantle
melting, Nature, 349, 283-284, 1991 (G.L.
Paper 2209).

Hemley, R. J., M. Hanfland, andH. K. Mao, High-
pressure dielectric measurements of solid hy-
drogen to 170 GPa, Nature, 350, 488-491,
1991 (G.L. Paper 2218).

Kubicki, J. D., G. E. Muncill, and A. C. Lasaga,
Chemical diffusion in melts on the
CaMgSi206-CaAl2Si2C>8 join under high pres-
sures, Geochim. Cosmochim. Acta, 54, 2709-
2715, 1990 (G.L. Paper 2199).

Kubicki, J. D., and A. C. Lasaga, Molecular dy-
namics and diffusion in silicate melts, in Dif-
fusion, Atomic Ordering, andMass Transport,
J. Ganguly, ed., pp. 1-50, Advances in Physi-
cal Geochemistry Series, Springer- Verlag,
New York, 1990 (G.L. Paper 2202; no reprints
available for distribution).

Kubicki, J. D., and A. C. Lasaga, Molecular dy-
namics simulations of pressure and tempera-
ture effects on MgSiC>3 and Mg2Si04 melts
and glasses, Phys. Chem. Minerals, 77, 661-
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Kudoh, Y., C. T. Prewitt, L. W. Finger,A.
Darovskikh, and E. Ito, Effect of iron on the
crystal structure of (Mg,Fe)SiC>3 perovskite,
Geophys. Res. Lett. , 17, 1481-1484, 1990
(G.L. Paper 2179).

Kushiro, I., and B. O. Mysen, Experimental stud-
ies of the system Mg2Si04-Si02-H2 at pres-
sures 10"2-10"10 bar and temperatures to

1650 C: Application to condensation and
vaporization processes in the primitive solar
nebula, in Progress in Metamorphic andMag-
matic Petrology, (D. S. Korzhinskiy Memo-
rial Volume), L. L. Perchuk, ed., Chapter 16,
pp. 4 1 1 -433 , Cambridge University Press, New
York, 1991 (G.L. Paper 2214).

Liu, X., and C. T. Prewitt, High- temperature dif-
fraction study of LnCoC>3 perovskites: A
high-order electronic phase transition. J. Phys.
Chem. Solids, 52, 441-448, 1991 (G.L. Paper
2204).

Mao, H. K., R. J. Hemley, and M. Hanfland,
Infrared reflectance measurements of the in-
sulator-metal transition in solid hydrogen,
Phys. Rev. Lett., 65, 484-487, 1990 (G.L.
Paper 2186).

Mao, H. K., Y. Wu, L. C. Chen, J. F. Shu, and A.
P. Jephcoat, Static compression of iron to 300
GPa and Feo.8Nio.2 alloy to 260 GPa: Implica-
tions for composition of the core, /. Geophys.
Res.,95, B13, 21,737-21,742, 1990 (G.L. Pa-
per 2194).

Mao, H. K., R. J. Hemley, Y. Fei, J. F. Shu, L. C
Chen, A. P. Jephcoat, Y. Wu, and W. A.
Bassett, Effect of pressure, temperature, and
composition on lattice parameters and density
of (Fe,Mg)Si03-perovskites to 30 GPa, /.
Geophys. Res.,96,B5,S069-S019, 1991 (G.L.
Paper 2212).

Mao, H. K., and R. J. Hemley, Optical transitions
in diamond at ultrahigh pressures, Nature,
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Mysen, B. O., Volatiles in magmatic liquids, in
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My sen, B. O., Relationships between silicate melt
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Mysen, B. O., Effect of pressure, temperature, and
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Mysen, B. O., Interaction between water and melt
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Nagahara, H. , I. Kushiro, and B . O. Mysen, Vapor-
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tion in the Interior of the Earth, S. Marumo,
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distribution).

Parise, J. B., Y. Wang, A. Yeganeh-Haeri, D. E.
Cox and Y. Fei, Crystal structure and thermal
expansion of (Mg,Fe)Si03 perovskite,
Geophys. Res. Lett., 17,2089-2092,\990(G.L.
Paper 2207; no reprints available for distribu-
tion).

Pickett, W. E., R. E. Cohen, and H. Krakauer,
Lattice instabilities, isotope effect, and high-
Tc superconductivity in La2-xBaxCu04)JP/ry1y.
Rev. Lett., 67, 228-231, 1991.

Ross, N. L., and R. M. Hazen, High-pressure
crystal chemistry of MgSi03 perovskite, Phys.
Chem. Minerals, 17, 228-237, 1990 (G.L. Pa-
per 2176).

Ross, N. L., J. F. Shu, R. M. Hazen, and T.
Gasparik, High-pressure crystal chemistry of
stishovite, Am. Mineral.,75, 739-747, 1990
(G.L. Paper 2185).

Ross, N. L., and K. Leinenweber, Single crystal
structure refinement of high-pressure ZnGe03
ilmenite, Z. Kristallogr., 191, 93-104, 1990
(G.L. Paper 2203; no reprints available for
distribution).

Stafford, T. W., Jr., P. E. Hare, L. Currie, A. J. T
Jull, and D. Donahue, Accuracy of North
American human skeleton ages, Quaternary
Research, 34, 111-120, 1990 (G.L. Paper

2175).

Stafford, T. W., Jr., P. E. Hare, L. Currie, A. J. T.
Jull, andD. J. Donahue, Accelerator radiocar-
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Stathoplos, L., and P. E. Hare, Amino acids in
planktonic foraminifera: Are they phyloge-
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Modern Aspects ofBiomineralization in Plants
and Animals, Proceedings of the Fifth Inter-
national Symposium on Biomineralization, R.
E.Crick,ed.,pp.329-338,PlenumPubl.Corp.,
New York, 1989 (G.L. Paper 2166; no reprints
available for distribution).

Velinsky, D. J., M. L. Fogel, J. F. Todd, and B. M.
Tebo, Isotopic fractionation of dissolved am-
monium at the oxygen-hydrogen sulfide inter-
face in anoxic waters, Geophys. Res. Lett., 18,
649-652, 1991 (G.L. Paper 2211).

Yoder, H. S., Jr., Heat transfer during partial
melting: An experimental study of a simple
binary silicate system,/. Volcanol. Geotherm.
Res. 43, 1-36, 1990 (G.L. Paper 2182).

Zeitler, P. K., B. Barreiro, C. P. Chamberlain, and
D. Rumble, IE, Ion-microprobe dating of zir-
con from quartz- graphite veins at the Bristol,
New Hampshire, metamorphic hot spot, Geol-
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reprints available for distribution).

Zhang, J., D. Ye, and C. T. Prewitt, Relationship
between the unit-cell volumes and cation radii
of isostructural compounds and the additivity
of the molecular volumes of carbonates, Am.
Mineral.,76, 100-105, 1991 (G.L. Paper 22 10).

Zheng, Z Z., D. X. Gu, Y, Xin, D. O. Pederson, L.
W. Finger, C G. Hadidiacos, andR. M. Hazen,
A new 1212-type phase: Cr-substituted
TlSr2CaCu207 with Tc up to about 110 K,
Modern Phys. Lett., 5, 635-642, 1991 (G. L.
Paper 2223; no reprints available for distribu-
tion).

GEOPHYSICAL LABORATORY

185

Personnel

July 1, 1990 to June 30, 1991

Research Staff

Charles T. Prewitt, Director
Francis R. Boyd, Jr.
Ronald E. Cohen1
Larry W. Finger
Marilyn L. Fogel
John D. Frantz
P. Edgar Hare
Robert M. Hazen
Russell J. Hemley
Thomas C. Hoering
T. Neil Irvine
Ho-Kwang Mao
Bjorn O. My sen
Douglas Rumble III
David Virgo
HattenS. Yoder, Jr.

Postdoctoral Associates

Zhaoxin Gong2
David Palmer3
Ellen K. Wright4

Research Associates

Jingzhu Hu
Jinfu Shu

Postdoctoral Fellows

John V. B adding5
Gray E. Bebout
James Brenan6
Yingwei Fei7

Michael Hanfland
David B. Joyce8
Paul L. Koch9
James D. Kubicki10
Charles Meade11
Craig M. Schiffries12
Hiroko Takahashi13
Willem L. Vos14
Jinmin Zhang15

Predoctoral Associate

Julie Kokis16

Research Interns

Craig Bates17
Jon Cramer18
Karen Durana19
Howard Lu20
Alistaire M. Moore21
Nicole Y. Morgan22

Supporting Staff

Andrew J. Antoszyk, Shop Foreman
Bobbie L. Brown, Instrument Maker
Stephen D. Coley, Sr., Instrument Maker
David J. George, Electronics Technician
Christos G. Hadidiacos,

Electronics Engineer
Marjorie E. Imlay, Assistant to the Director
Lavonne Lela, Librarian23
Yunye Luo, Library Technician24
Harvey J. Lutz,

Technician/Mail Supervisor

186

CARNEGIE INSTITUTION

Mary M. Moore,

Word Processor Operator

— Receptionist
Lawrence B. Patrick,

Maintenance Supervisor25
David Ratliff, Jr.,

Maintenance Technician26
Pedro J. Roa,

Maintenance Technician27
Susan A. Schmidt,

Coordinating Secretary
John M. Straub,

Business Manager
Mark Vergnetti,

Instrument Maker28
Stephanie Vogelpohl,

Administrative Assistant29

J. Michael Palin, Yale University
Nicolai P. Pokhilenko, Inst. Mineralogy &

Petrology, Novosibirsk, USSR
Robert Popp, Texas A. and M.
Guoyin Shen,

University of Uppsala, Sweden
Bradley Tebo,

Scripps Institution of Oceanography
Noreen C. Tuross, Smithsonian Institution
K. Vedam, Pennsylvania State University
David von Endt, Smithsonian Institution
YanWu,

University of California, Berkeley

Adjunct Senior Research Scientist
Peter M. Bell

Visiting Investigators

Emeritus

Rateb M. Abu-Eid,

Kuwait Institute for Scientific Research
Constance Bertka,

Arizona State University
Alison Brooks,

George Washington University
Robert T. Downs,

Virginia Polytechnic Institute

& State University
Glenn A. Goodfriend,

Weizmann Institute of Science, Israel
Matthew Hoch, University of Delaware
Hans G. Huckenholz,

Munich University, Germnay
Donald G. Isaak,

Naval Research Laboratory
James G. Kirklin,

Johns Hopkins University
Kevin Mandernack,

Scripps Institution of Oceanography

Hatten S. Yoder, Jr., Director Emeritus
Felix Chayes, Petrologist Emeritus

Appointed Sept. 1, 199U

2 Appointed Jan. 28, 1991

3 Appointed Oct. 1, 1990

4 To June 30, 1991

5 Accepted position as Assistant Professor, The
Pennsylvania State University

6 Appointed Nov. 15, 1990

7 Accepted position as Associate Staff Member,
Geophysical Laboratory

8 To June 30, 1991

9 Appointed Sept. 1, 1990
°To December 30, 1990

1 Appointed Gilbert Fellow July 1, 1990
2To Sept. 30, 1990

3 Appointed March 1, 1991

4 Appointed April 1, 1991
5To June 30, 1991
Appointed July 1, 1990
7FromJune24, 1991
8FromMay 15, 1991
9FromJune 1, 1991

20From June 24, 1991
21 From June 24, 1991
22From June 24, 1991

GEOPHYSICAL LABORATORY 187

23 Also associated with the Department of
Terrestrial Magnetism (DTM)

24 Also associated with DTM

25 Also associated with DTM

26 Also associated with DTM

27 Also associated with DTM

28 To Nov. 30, 1990

29AppointedJune3, 1991

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