JOURNAL OF PETROLOGY
VOLUME 48
NUMBER 7
PAGES 1369^1386
2007
doi:10.1093/petrology/egm022
Crystallization and Degassing in the Basement
Sill, McMurdo DryValleys, Antarctica
ALAN BOUDREAU1* AND ADAM SIMON2
1
DIVISION OF EARTH AND OCEAN SCIENCES, NICHOLAS SCHOOL OF THE ENVIRONMENT AND EARTH SCIENCES,
DUKE UNIVERSITY, DURHAM, NC 27708, USA
2
DEPARTMENT OF GEOSCIENCE, UNIVERSITY OF NEVADA, LAS VEGAS, NV 89154-4010, USA
RECEIVED AUGUST 8, 2006; ACCEPTED MARCH 30, 2007
ADVANCE ACCESS PUBLICATION MAY 23, 2007
The Basement Sill is part of the Ferrar Large Igneous Province
exposed in the McMurdo Dry Valleys, Antarctica. The sill is
330 m thick in the Bull Pass area and 450 þ m thick in the Dais
area, 12 km to the west, and is characterized by phenocryst-free
lower and upper margins and an orthopyroxene-rich central ‘tongue’
(opx 1^5 mm in size). Halogen variations in apatite from a suite of
samples collected along vertical transects through the sill were examined to evaluate the process of crystallization-induced degassing
(i.e. second boiling) and its effects on magma chemistry. Apatite
grains from any given sample are generally unzoned with respect to
Cl and F concentrations, but may vary by 20^30 mol% in the halogen site between grains. Overall average Cl/F mass ratios increase
with height from the lower margin to the center of the sill, and then
decrease to near zero towards the top margin where the rocks are
relatively oxide-rich. The Cl/F trend parallels those of bulk MgO
and grain size. The upper margin contains abundant mafic pegmatoids and the apatite in these segregations has lower Cl/F ratios
compared with that in the host-rocks, although REE show no
measurable difference. Numerical modeling illustrates that a cooling
and crystallizing sill initially develops two separate vapor-saturated
zones at the lower and upper margins owing to the irreversible heat
loss to the cooler country rock. Vapor separating from the lower zone
migrates upward into hotter silicate liquid, where it is resorbed, thus
increasing the Cl/F mass ratio of the liquid. This process leads to
saturation and precipitation of apatite from the liquid with a higher
Cl/F ratio than would otherwise occur. Volatile enrichment can also
aid compaction and grain growth in the central part of the sill.
In contrast, the relatively Fe-rich, Cl-poor nature of the upper zone
rocks suggests that these rocks may have crystallized from more
evolved, degassed silicate liquid, possibly compacted out of the underlying crystal mush. In addition, as vapor sourced from the lower and
central parts of the sill ascends into the cooler upper zone of the sill,
*Corresponding author. Telephone: 919-684-5646. Fax: 919-684-5833.
E-mail: [email protected]
the vapor may be localized (along with late interstitial silicate
liquid) to form pegmatoids at temperatures at which Cl is less
favored in apatite and can be leached from existing apatite by the
ascending vapor, the latter causing the observed decrease in the Cl/F
mass ratio of apatite in the (evolved) pegmatoids.
KEY WORDS:
Ferrar Igneous Province; halogens; fluid; apatite
I N T RO D U C T I O N
Magmatic volatile phases can play a number of important
roles in the generation, crystallization, and eruption of mafic
magmas (Roggensack et al., 1997; Huppert & Woods, 2002;
Yang & Scott, 2002; Cervantes & Wallace, 2003; Wallace,
2003; Cashman, 2004; Gonnerman & Manga, 2005;
Hammer, 2006). However, in plutonic rocks the evidence of
degassing is difficult to detect given the fugitive nature of
volatile fluids (Candela & Blevin, 1995). Halogen variations
in apatite are one mechanism that has been found useful in
deciphering the complex degassing history of layered intrusions and associated mafic rocks (e.g. Brown & Peckett, 1977;
Boudreau & McCallum, 1989; Willmore et al., 2000).
This study details the halogen composition of apatite in
dolerite from the Basement Sill from the west Bull Pass and
Dais areas (where the sill is 330^350 and 450þ m thick,
respectively; Marsh, 2004) of the Ferrar Igneous Province
in the McMurdo Dry Valleys, Antarctica. The crystallized
magma bodies provide an excellent field laboratory for the
study of vapor migration in a solidifying crystal pile as
both were emplaced as an orthopyroxene-phyric mush
(Marsh, 2004) with a relatively simple emplacement and
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JOURNAL OF PETROLOGY
VOLUME 48
crystallization history that can otherwise complicate the
understanding of the compositional evolution of vapor
from a basic igneous intrusion.
G E N E R A L G EOLO GY A N D
P R E V I O U S I N T E R P R E TAT I O N S
The Ferrar Large Igneous Province is located along strike
of the Trans-Antarctic Mountains, which run north^south
and divide Antarctica into the Eastern and Western ice
sheets (Gunn & Warren, 1962; McKelvey & Webb, 1962;
Gunn, 1966; Borg et al., 1990; Elliot & Fleming, 2004).
The province was produced during the Jurassic-age
(Heimann et al., 1994; Knight & Renne, 2005) fragmentation of Gondwanaland. In the McMurdo Dry Valleys in
Southern Victoria Land the Ferrar consists of a series of
four vertically interconnected dolerite sills (Fig. 1), interconnecting dykes and extrusive rocks (i.e. the 500 m
thick Kirkpatrick flood basalt province). The high-aspect
ratio dolerite sills vary in thickness from 100 to 500 m and
crop out over an area of the order of 10 000 km2. In order
from bottom to top they are called the Basement Sill,
Peneplain Sill, Asgard Sill and Mt. Fleming Sill (Gunn &
Warren, 1962; Hamilton, 1965). The sills are typically
separated from one another by several hundred meters,
except in the area of Bull Pass, where the Basement Sill
and Peneplain Sill are in contact with one another and
were interpreted by Marsh (2004), on the basis of field
evidence, to have formed from the top down.
The compositions of the Ferrar sills are dominantly
quartz-normative dolerite. The lowermost exposed sill, the
Basement Sill, is composed of a lower phenocryst-free
gabbronorite chilled marginal zone, an orthopyroxene
phenocryst-bearing central zone, identified as the
‘Opx tongue’ by Marsh (2004), and a phenocryst-free
gabbronorite upper margin (Gunn & Warren, 1962;
Hamilton, 1965; Marsh, 2004). In the West Bull Pass area,
the lower chilled margin is dark, aphanitic to microcrystalline, contains plagioclase, orthopyroxene, augite and
inverted pigeonite, and has an abrupt contact with the
overlying opx tongue. Moving up from the lower margin,
the grain size and opx content of the sill both increase
toward the center of the Opx tongue, and then both
decrease toward the top margin. The increase in orthopyroxene modal abundance in the center of the sill manifests
itself in a strong change in bulk-rock MgO concentration
from 7 wt % in both the upper and lower margins to
20 wt % in the tongue (Marsh, 2004). The dark-colored
upper margin superficially mimics the lower chilled
margin and is comprised dominantly of plagioclase, orthopyroxene, augite and inverted pigeonite; however, the
upper margin has a more Fe-rich bulk composition and,
for the most part, comparatively larger grain sizes (fineto medium grained) relative to the aphanitic lower chilled
NUMBER 7
JULY 2007
margin. The difference in grain size between the lower and
upper chilled margins probably reflects higher country
rock temperatures above the sill at the time of emplacement owing to the high magma flux feeding the
Kirkpatrick flood basalts and the formation of earlier sills
(i.e. Peneplain Sill) above the Basement Sill (Hersum et al.,
submitted). Oxides (magnetite and ilmenite) are particularly abundant at the very top of the upper margin.
Another difference between the lower and upper margins
is the common presence of mafic pegmatoids and granophyric segregations beginning in the upper part of the
opx tongue and continuing to the top of the sill (Zavala &
Marsh, 2001, 2002). The pegmatoid segregations are
composed dominantly of plagioclase and pyroxene (the
latter to several centimeters in length) with minor modal
amounts of quartz, alkali feldspar, amphibole, apatite, biotite and oxides. The pegmatoids are enriched in SiO2, FeO
and incompatible elements (e.g. K, Ba, Rb, Nb and Ti)
and have higher Ab:An (25%) ratios in plagioclase
relative to plagioclase in the host-rock. The liquids that
crystallized to yield the pegmatoids are inferred to
represent the end products of the crystallization of latestate interstitial liquids that were probably expelled
(i.e. filter pressed) from the surrounding gabbronorite
during compaction of the magma (Geist et al., 2005).
The abrupt contact between the lower and upper
margins and the Opx tongue manifests itself in a distinct
break in slope noted in a crystal size distribution study
(Zieg & Forsha, 2005) as well as a dramatic rise in
temperature (2008C) inferred from two-pyroxene
thermometry (Simon & Marsh, 2005). These results reflect
directly the presence of the extensive Opx tongue and
suggest that the pyroxene phenocrysts in the tongue were
entrained in the ascending magma after having been
texturally equilibrated at a much deeper level of the
magmatic plumbing system than the present level of sill
emplacement. Notably higher concentrations of Cr and Al
in the Opx phenocrysts in the tongue relative to the upper
and lower margins lend additional support to this hypothesis (Simon, in preparation). The magma that formed most
of the Basement Sill probably intruded as a crystal-rich
mush (i.e. silicate liquid þ phenocrysts; Marsh, 2004;
Charrier & Marsh, 2005; Petford et al., 2005) and the
spatial position of the tongue reflects the inability of
the large orthopyroxene phenocrysts to settle through
the viscous lower margin.
Finally, the Dais Intrusion makes up the Dais, a
topographic feature about 12 km west of the Bull pass
area, and is an extension of the Basement Sill. It differs
from the sill in the Bull Pass area in being thicker
(the exposed portion is 450 m thick; the lower contact
is not exposed) and in having more pronounced modal
layering, particularly in the upper part of the Opx tongue
(Marsh, 2004). The layering is defined by sharp changes
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DEGASSING IN BASEMENT SILL, ANTARCTICA
Fig. 1. Top: geological map of part of the DryValleys, Antarctica, showing locations of the Bull Pass and Dais areas (after Marsh & Zieg, 2004).
The Asgard and Mt. Fleming sills are spatially closely associated and are shown here as one unit. Bottom: simplified sections of the Basement
Sill in the Dais and west Bull Pass areas.
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JOURNAL OF PETROLOGY
VOLUME 48
in the amounts of orthopyroxene and plagioclase and
ranges in thickness from centimeter to several meter scale.
Petrology of late-crystallizing minerals
All halogen-bearing minerals are present in only minor
modal abundance and none appear to have been early
liquidus (i.e.cumulus) phases. Apatite is the principal
mineral of interest as it is the most commonly occurring
(and probably the first precipitating) halogen-bearing
mineral. Neither mica nor amphibole shows any textural
evidence of crystallizing before apatite. This is consistent
with evidence from layered intrusions crystallized in
a low-pressure environment (such as the Skaergaard
Intrusion) where apatite is the first halogen-bearing
‘cumulus’ mineral observed after 90% crystallization
(e.g. Wager & Brown, 1967). The crystallization sequence
of a representative chilled margin composition was modeled using MELTS (Ghiorso & Sack, 1995), which predicts
that apatite is the sole halogen-bearing mineral during
cooling. Also, unlike the micas and amphiboles, which
have a strong crystal-chemical control on halogen substitution, the apatite halogen solid solution (i.e. Cl , F) is ideal
at high temperatures (e.g. Volfinger et al., 1985; Tacker &
Stormer, 1989).
The apatite grains analyzed in this study are interstitial
to early formed minerals (plagioclase and pyroxene),
and are typically associated with other late-crystallized
phases such as quartz, Fe^Ti oxides and, relatively rarely,
biotite. Apatite modal abundance is qualitatively correlated with (unpublished) whole-rock P2O5 concentrations,
suggesting that apatite is the dominant phosphorusbearing phase in the rocks. In most instances apatite
displays a characteristic acicular habit, forming thin rods
NUMBER 7
JULY 2007
usually no more than a few tens of microns in width but
locally up to about 300 mm in length. In the central part
of the sill, however, grains can be considerably thicker
and have a tabular habit (Fig. 2).
Finally, apatite from a few samples of the exposed part
of the Dais was analyzed to see if lateral compositional
variation is present in the Basement Sill between the Bull
Pass and Dais Intrusion.
M ET HODS
Apatite grains in polished thin sections were analyzed by
wavelength-dispersive spectrometry (WDS) using the
Cameca Camebax electron probe microanalyzer (EPMA)
at Duke University, Durham, North Carolina, USA.
Typical analytical conditions were: 15 keV acceleration
voltage, 15 nA beam current, and a focused beam
(a requirement given the typically small size of individual
grains). Peak and background counting times for Cl and F
were in the ranged of 20^30 s and 10^15 s, respectively.
Other elements were counted on the peak for 30 s and
background for 15 s. To minimize volatilization, F and Cl
were counted first. Standards included natural chlorapatite
(RM-1, Morton & Catanzaro, 1964) and fluorapatite
(from Wilberforce, Ontario) for Ca, P and the halogens,
and allanite (102522, #42; Frondel, 1964) for La and Ce
for which the analyte lines are relatively free from interferences (Roeder, 1985). Standard intensities for F and Cl
using the apatite standards were also checked against
fluorite, topaz and halite. Fluorine was analyzed using
a synthetic Si^W layered diffracting crystal. The OH ion
and H2O were calculated by difference based on ideal
OH site occupancy. Previous work with these analytical
Fig. 2. Backscattered electron image of apatite from the basal Ferrar sill, west Bull Pass, Antarctica. Sample 31U: acicular apatite from the lower
chilled margin, just below the orthopyroxene tongue. (See Table 1, analysis 1, for analysis.) Sample 14N: large, tabular apatite from the center of
the orthopyroxene tongue. (See Table 1, analyses 5 and 6, for core and rim analysis.) Ap, apatite; Pl, plagioclase; Opx, orthopyroxene; Q, quartz.
(Note different scales.)
1372
BOUDREAU & SIMON
DEGASSING IN BASEMENT SILL, ANTARCTICA
conditions (see Willmore et al., 2000) has shown no significant apatite orientation effects on the analyses as
reported by Stormer et al. (1993). However, to minimize
any potential orientation effects, preference was given to
grains cut parallel to the c-axis.
R E S U LT S
Apatite composition
Representative apatite compositions are listed in Table 1
(a complete set of analyses is available as an Electronic
Appendix at http://www.petrology.oxfordjournals.org/)
and those from west Bull Pass are shown as a function of
height in the sill in Fig. 3. The mole fraction of F ranges
from 04 to end-member fluorapatite (and a few analyses
for which the mole fraction F exceeds 10), whereas Cl
ranges from nears its limit of detection (LOD 001
wt %) to a maximum of about 05 mole fraction. As has
now been observed in a number of studies (e.g. Boudreau
et al., 1986, 1992, 1995; Willmore et al., 2000), individual
apatite grains from any given sample tend to be homogeneous and not significantly zoned in Cl and F (e.g.
Table 1, analysis 5 and 6), whereas different grains in the
same sample can vary by as much as 30þ mol% in the
hydroxyl/halogen site. Thus, detailed stratigraphic changes
can be masked by this within-sample variation. However,
overall there remains a broad increase in Cl and decrease
in F from the lower margin towards the center of the sill,
and Cl then decreases to below LOD in the upper margin.
These stratigraphic changes in Cl broadly parallel trends
in the bulk-rock MgO and grain size (Fig. 3).
The compositions of Ferrar apatite along with compositional fields from a number of other mafic intrusions are
plotted in an F^Cl^OH ternary diagram in Fig. 4.
The few analyses from the Dais overlap those from the
Bull Pass area, suggesting that there are no significant
regional variations in apatite chemistry. The Cl-poor
nature of most Ferrar apatite is typical of intrusions such
as the Skaergaard (Greenland), Great Dyke (Zimbabwe),
Kiglapait (Canada), and Munni Munni (Australia),
where the mole fraction of the chlorapatite component is
generally less than 20 mol% (see Boudreau, 1995). Except
for the few Cl-rich samples from the central part of the
Opx tongue, none are as Cl-rich as is seen in the lower portions of the Stillwater and Bushveld intrusions below the
economically important platinum-group element zones of
Table 1: Selected electron microprobe analyses of apatite in the basal Ferrar Sill, Antarctica
Analysis:
1
2
3
4
5
6
8
9
10
11
Sample:
31U
27N
24N
17U
14N-core
14N-rim
06N
01N
DG-101
DG-103
Location:
Bull Pass
Bull Pass
Bull Pass
Bull Pass
Bull Pass
Bull Pass
Bull Pass
Bull Pass
Dais
Dais
Height (m):
3
25
83
163
196
196
284
339
near top
near top
Rock type:
gabbro
opx’nite
opx’nite
opx’nite
opx’nite
opx’nite
gabbro
gabbro
pegmatoid
host gabbro
wt %
CaO
553
551
544
555
552
544
547
547
545
554 (05)
P2O5
411
409
414
410
410
411
418
410
414
406 (04)
F
275
269
249
159
241
221
242
362
345
Cl
048
077
088
173
057
058
024
012
018
329 (024)
030 (003)
H2Oy
035
030
034
059
049
056
055
001
007
014
La3O3
n.a.
n.a.
n.a.
n.a.
n.a.
n.a.
n.a.
n.a.
010
014 (005)
Ce2O3
n.a.
n.a.
n.a.
n.a.
n.a.
n.a.
n.a.
n.a.
044
042 (005)
SiO2
n.a.
n.a.
n.a.
n.a.
n.a.
n.a.
n.a.
n.a.
015
029 (002)
Na2O
n.a.
n.a.
n.a.
n.a.
n.a.
n.a.
n.a.
n.a.
013
005 (002)
Total
O F,Cl
Total
999
126
986
1006
131
993
995
125
983
1004
106
994
989
105
978
998
114
986
997
107
987
994
155
978
1004
149
989
1006
145
992
XCl
007
011
013
025
08
08
003
002
003
004
XF
073
072
068
042
64
60
065
098
093
088
XOHy
020
017
020
033
28
32
031
001
004
008
n.a., not analyzed; opx’nite, orthopyroxenite. Numbers in parentheses for analysis 11 are 1s counting statistic errors and
are typical for all analyses.
Height relative to lower contact.
y
H2O and OH calculated on the bases of Cl þ F þ OH ¼ 1 per formula unit.
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Fig. 3. Summary results from the Ferrar Sill, Bull Pass area. (a) Mole fraction Cl (filled symbols and F (open symbols) in apatite; (b) molar
Cl/F ratio of apatite; (c) bulk-rock MgO concentration (wt %); (d) average orthopyroxene grain size.
these two intrusions. Apatite in the Stillwater and Bushveld
intrusions is strongly enriched in the end-member chlorapatite component throughout rather sizeable stratigraphic
thicknesses.
Variations in apatite compositions across
a pegmatoid^host-rock contact
Apatite was analyzed in three samples collected across
a pegmatoid^host-rock contact from the Dais layered
intrusion: one from within the pegmatoid segregation,
one from the pegmatoid^gabbronorite host-rock contact
and one 3 m into the host-rock. The samples were collected
from the upper part of the Dais intrusion stratigraphically
above the Opx tongue. Data from these apatite analyses
are shown in Fig. 5 and example analyses are provided in
Table 1; analysis 10 is apatite from within the pegmatoid
and 11 is apatite from the host gabbro. Apatite within the
pegmatoid is Cl poor (all 502 wt % Cl) and most grains
are nearly end-member fluorapatite. Furthermore, apatite
in the pegmatoid contains even less Cl than apatite in
the surrounding upper zone host-rock. Bulk analysis
of the pegmatoid demonstrates that the concentrations of
incompatible trace elements are elevated between two
and three times their concentrations in the host-rock
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DEGASSING IN BASEMENT SILL, ANTARCTICA
Fig. 4. Comparison of Ferrar apatite from the Bull pass area (*) and the Dais Intrusion (þ) with those from the Stillwater and Bushveld
complexes (only the range below the platinum reefs is shown), as well as the Munni Munni and Great Dyke. Inset shows the range of apatite
from associated sill, dike and marginal rocks associated with these other intrusions. After Boudreau (1995).
(Geist et al., 2005); however, the concentrations of rare
earth elements (REE) in apatite are essentially unchanged
from host gabbronorite to the pegmatoid.
DISCUSSION
To summarize, apatite in the Basement Sill shows significant Cl/F variations both within individual samples and
as a function of stratigraphic position, with the highest
Cl/F ratios occurring in the middle of the sill. Assuming
that the apparent rapid quench typical of the lower chilled
margin preserves apatite halogen compositions in equilibrium with the initial (unfractionated) liquid, the initial
apatite is characterized by X(Cl) 005 and X(F) 07.
Despite the complexity of the magma dynamics indicated by the characteristics of the Opx tongue, a notable
feature of the sill is that, once emplaced, it cooled uniformly from the top and bottom such that the silicate
liquid would have become vapor-saturated nearly simultaneously at both margins. The sill took some 1000 years
to solidify and evidence for post-solidification lowtemperature hydrothermal alteration is absent (Simon &
Marsh, 2005). As mentioned above, the finer grain size in
the lower margin relative to the upper margin suggests
that the upper margin may have cooled slightly more
slowly; the proximity of the Peneplain Sill above the
Basement Sill may have played a role in this (see Hersum
et al., submitted). Thus, the rapid crystallization time and
preservation of magmatic halogen ratios in trace phases
such as apatite provide the opportunity to study the loss
of vapor evolved from the upper margin to the overlying
country rock and the migration of vapor evolved from the
lower margin as this vapor ascended through the sill.
Prior to the interstitial liquid becoming saturated in
a halogen-bearing phase (e.g. apatite, biotite, amphibole
or vapor), one would not expect there to be fractionation
of halogens by crystallization of the observed anhydrous
minerals. Ignoring vapor for the moment, and although it
appears that apatite is the first halogen-bearing mineral
phase to crystallize, the crystallization of apatite, biotite
or amphibole should all incorporate F in preference to Cl
(e.g. Speer, 1984; Candela, 1986), particularly in rocks with
high Mg/Fe ratio (e.g. Volfinger et al., 1985). Considering
only apatite, Beswick & Carmichael (1978) noted that
fluorapatite is generally more stable (i.e. crystallizes earlier) than chlorapatite and hydroxyapatite at magmatic
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Fig. 5. Comparison of apatite traverse from host-rock (upper left photomicrograph), through contact (center photomicrograph) and to
pegmatoid (upper right photomicrograph) from the upper margin of the Dais Intrusion. Plots are of (a) Cl, (b) F, and (c) Ce2O3 concentration
(wt %) in apatite as a function of relative distance across the thin section shown in the photomicrographs at the top.
temperatures and the phase equilibrium interpretations of
Tacker & Stormer (1993) are consistent with their observations. This suggests that apatite behaves similarly to other
halogen- and OH-bearing minerals in that F-bearing
end-members have significantly higher thermal stabilities
(Patin‹o Douce & Roden, 2006). Hence, Cl/F variations in
apatite are analogous to Mg/Fe variations in olivine or
orthopyroxene (Cawthorn, 1994), such that an increase
in the Cl/F ratio of apatite implies both an increase in the
Cl/F ratio of the interstitial silicate liquid and that Cl-rich
apatite formed from the primary Basement Sill silicate
liquid should be stable only at lower temperatures. Thus,
the Cl/F mass ratio in the silicate liquid should increase
after apatite and other Cl- and F-sequestering mineral
phases begin to crystallize from the interstitial liquid.
An exsolving vapor will effectively decrease the Cl/F
mass ratio of the silicate liquid as degassing progresses
(as discussed more fully below). As pockets of interstitial
silicate liquid that are saturated in apatite or other halogen-bearing phases become isolated during the final
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DEGASSING IN BASEMENT SILL, ANTARCTICA
crystallization of interstitial liquid, apatite of different
compositions can be produced. This is considered to be
the main cause of the observed compositional variation
in apatite within individual samples.
The overall stratigraphic changes in the Cl/F and OH/F
ratio of apatite can be caused by several crystallization
mechanisms, some of which have been discussed previously
(Boudreau et al., 1986; Boudreau & McCallum, 1989;
Boudreau & Kruger, 1990; Mathez & Webster, 2005).
The effect of temperature and of variations in the activities
of F and Cl on apatite compositions have been described
by Boudreau et al. (1992), who considered the equilibrium
between apatite and an aqueous fluid, for which thermodynamic data are available, that are both in equilibrium
with a silicate liquid. Although the activities of the halogens in a silicate liquid are not expressed explicitly by
such a treatment, an increase in the activity or chemical
potential of either of the halogens in any phase must
occur in all other coexisting halogen-bearing phases
as well.
Let us consider the exchange reaction between apatite
(ap) and an aqueous fluid (af) in which HCl and HF are
present as neutral species:
F apatitesolid þ HClaf ¼ Cl apatitesolid þ HFaf ð1Þ
and the accompanying apparent exchange constant:
ap af XCl aHF
log K ¼ log ap
:
XF aaf
HCl
ð2Þ
j
In equation (2), Xi denotes the mole fraction of component
j
i in phase j and ai denotes the activity of i in phase j.
Activity coefficients for these species are not known and,
thus, are not included in the apparent exchange constant.
The relationship between fluid composition and apatite
composition is given by
ap af XCl
aHCl
:
ð3Þ
ap ¼ K
XF
aaf
HF
Analogous equations can be written to describe OH^F
exchange between apatite and aqueous fluid. Using the
observation that halogen substitution in apatite is ideal
above 5008C (Tacker & Stormer, 1989), Boudreau et al.
(1992) fitted expressions for log K to a simple temperature-dependent function. Shown in Fig. 6 are calculated
apatite compositions that result from changing either
the Cl/F or OH/F ratio of the silicate liquid at a fixed temperature of 11008C (dashed lines) or from changing the
temperature at fixed HCl/HF and H2O/HF activity ratios
of 30 and 100, respectively (continuous lines). The model
trends in Fig. 6 suggest that the observed variations in
apatite compositions may be caused by changing either
the HCl/HF and H2O/HF ratios by approximately an
order of magnitude, or by varying the temperature
at which apatite equilibrates by up to 5008C.
As noted above, there do not appear to be any halogensequestering minerals crystallizing in the sill prior to
apatite saturation; hence, fractional crystallization of the
observed halogen-free minerals alone cannot lead to variations in halogen ratios as a function of stratigraphic
position. That is, because P behaved as an incompatible
trace element during the crystallization of the Basement
Sill, and assuming as a first approximation that apatite
crystallization was controlled only by the P concentration
in the silicate liquid, then all interstitial silicate liquid
should have reached apatite saturation after the same
amount of crystallization. The system would have followed
a single liquid line of descent with respect to apatite
crystallization and apatite would have begun to crystallize
at about the same temperature throughout the intrusion.
Because this is inconsistent with a large temperature
range in apatite equilibration, temperature effects alone
cannot be the major cause of the stratigraphic variation in
apatite composition. One would also expect the same for
biotite and amphibole crystallization.
However, the separation of an aqueous fluid phase from
a crystallizing silicate liquid can fractionate the halogens
owing to the strong affinity of Cl relative to F for an
exsolved aqueous fluid phase relative to silicate liquid
(Sourirajan & Kennedy, 1962; Helgeson, 1964; Burnham,
1967; Roedder, 1984, 1992; Bodnar et al., 1985; Candela,
1986; Chou, 1987; Shinohara et al., 1989; Chou et al., 1992;
Metrich & Rutherford, 1992; Sterner et al., 1992; Anderko
& Pitzer, 1993; Cline & Bodnar, 1994; Lowenstern, 1994;
Piccoli & Candela, 1994; Webster et al., 1999, 2004;
Webster, 2004; Webster & DeVivo, 2004; Mathez &
Webster, 2005). Furthermore, as discussed below, migration
of exsolved vapor through a crystallizing magma can
readily redistribute halogens from one region to another,
leading to variations in halogen ratios in both the silicate
liquid and apatite.
We suggest that vapor evolved from the lower half of the
sill and subsequently migrated upward through the crystal
mush where it encountered hotter, fluid-undersaturated
silicate liquid (i.e. silicate liquid that had not reached
crystallization-induced fluid saturation). The free vapor
phase would have had a high Cl/F mass ratio owing to the
strong mass transfer of Cl from silicate liquid to vapor
relative to F (Dingwell & Scarfe, 1983; Webster &
Holloway, 1990; Candela & Piccoli, 1995). As the vapor
ascended into hotter silicate liquid, the vapor would have
been resorbed by the silicate liquid; hence, resulting in elevated volatile concentrations and higher Cl/F ratios in
the interstitial silicate liquid in the central part of the sill.
This process would ultimately yield higher Cl/F ratios in
apatite crystallized from this interstitial silicate liquid.
In contrast, early formed apatite crystallized from the
upper margin of the Basement Sill and its associated
pegmatoidal and granophyric segregations grew from
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Fig. 6. Plots of (a) the Cl/F ratio as a function of height and (b) OH/F ratio as a function of height, both from the Basal Sill from the Bull Pass
area. In (a), the continuous lines show the calculated Cl/F ratio of apatite in equilibrium with an aqueous fluid with a fixed HCl/HF activity
ratio of 30 and at the various temperatures indicated. The dashed lines are the calculated Cl/F ratio of apatite in equilibrium with an aqueous
fluid at 11008C and with different HCl/HF activity ratios as noted. In (b) the continuous lines show the calculated OH/F ratio of apatite in
equilibrium with an aqueous fluid with a fixed H2O/HF activity ratio of 100 at the various temperatures indicated. The dashed lines are the
calculated OH/Fratio of apatite in equilibrium with an aqueous fluid at 11008C and with different H2O/HFactivity ratios indicated. (See text for
additional discussion.)
more evolved silicate liquid that had already degassed and
preferentially lost an unquantified amount of Cl to the
exsolved volatile phase. In addition, aqueous fluids separating from the lower and eventually from the middle of
the intrusion migrated into cooler rocks of the upper
margin. For aqueous fluids with a fixed HCl/HF ratio,
the fluids would be in equilibrium with lower Cl/F apatite
and could leach Cl from pre-existing apatite.
Evidence from pegmatoids
The strong enrichments in SiO2, FeO and incompatible
trace elements in the pegmatoid (Geist et al., 2005) are
consistent with these segregations having crystallized
from late-stage, evolved interstitial silicate liquid. The lack
of variation in rare earth element (REE) apatite noted
in this study suggests that the higher bulk-rock REE
concentrations in the pegmatoids simply reflect a higher
amount of this evolved liquid than is seen in the host-rock,
but that the evolved liquid in both the pegmatoid and the
interstitial liquid of the host otherwise have similar REE
concentrations. However, the lower Cl/F ratio of apatite
in the pegmatoid as compared with host apatite implies
that the apatite equilibrated with a liquid or vapor with
lower Cl/Fratio or that this ratio was fixed but equilibrated
at a lower temperature, or both.
It is suggested that the pegmatoidal texture developed
as vapor, evolved from the lower parts of the sill,
used these regions of higher amounts of silicate liquid
(i.e. pegmatoids) as preferential conduits through which
to ascend. Further, these aqueous fluids were relatively
Cl-poor and were not effective at moving significant
quantities of REE. Data from saline aqueous inclusions
in quartz hosted by granophyre of the Capitan Pluton
(New Mexico) suggest that the light REE (LREE) may
be fractionated from the heavy REE (HREE) during
volatile phase exsolution, with the LREE having been
scavenged by the saline aqueous fluid (Banks et al., 1994).
In another study, whole-rock analyses of potassically
altered and unaltered granodiorite also showed enrichment in LREE in the aqueous-fluid-altered rocks (Taylor
& Fryer, 1980). The ability of aqueous fluids to scavenge
preferentially the REE from crystallizing silicate liquid
has also been shown experimentally (Flynn & Burnham,
1978; Sakagawa, 1989; Reed, 1995; Reed et al., 2000).
Although the mass transfer of individual REE from silicate
liquid to vapor is a complex function of the total salinity of
the vapor phase (Candela, 1990) and the concentrations of
all other chloride-bound elements in the vapor (Reed, 1995;
Reed et al., 2000), in general, the exsolution of a Cl-bearing
aqueous fluid from a silicate liquid will lead to a decrease
in the LREE/HREE mass ratio of the silicate liquid.
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DEGASSING IN BASEMENT SILL, ANTARCTICA
The partition coefficients describing the simple mass transfer of La and Lu between silicate liquid and vapor increase
from approximately 001 and 0001 when the Cl concentration of the vapor is 001molal to 2 and 07 when the Cl
concentration of the vapor is 2 molal. Thus, if ascending
vapor is resorbed into stratigraphically higher, hotter
silicate liquids (as discussed below), this should increase
the LREE concentration of that interstitial silicate liquid
and, even upon subsequent vapor saturation of that parcel
of silicate liquid, the LREE/HREE mass ratio might be
expected to be higher than that of silicate liquid stratigraphically lower in the magma chamber. The fact that
REE concentrations in the pegmatoid and host-rock
apatite are equivalent suggests that the late-state interstitial silicate liquid that crystallized to form the pegmatoids
did not resorb a significant quantity of ascending vapor
and, further, that the vapor exsolving from this segregated
silicate liquid was of a low enough salinity or temperature
that it did not effectively fractionate the REE.
Modeling degassing in a cooling
and crystallizing intrusion
Vapor separation in the cooling and crystallizing Basement
Sill was modeled using a modified version of the
IRIDIUM program (Boudreau, 2003). This program uses
a simplified version of the MELTS free energy minimization program (Ghiorso et al., 1994; Ghiorso & Sack, 1995)
coupled with one-dimensional heat and mass transfer to
model a number of igneous processes such as compaction
and silicate liquid percolation through a porous assemblage. The program has been modified to include a more
complete C^O^H^S fluid (see the Appendix). It also has
been modified to allow for an upper and lower thickness
of material to act as a thermal sink; originally the program
simply assumed a fixed heat loss from the top and bottom
of the magma column. This modification allows
more realistic modeling of the thermal profile of a sill
that is losing heat by thermal conduction into cooler
country rock.
The initial conditions of the model assume a 350 m
thick, pyroxene-phyric sill that is allowed to cool and
crystallize as heat is lost to the over- and underlying country rocks. Because we are interested specifically in the
effects that volatile saturation and volatile migration may
have on halogen ratios, the model initially assumes a
uniform mush column. (The phenocryst-free marginal
rocks are ignored as they probably represent an initial
phenocryst-free leading edge of intruding magma that
was essentially chilled prior to the emplacement of the
Opx tongue.) Paleodepth estimates suggest country rock
temperatures as low as 1508C (Simon & Marsh, 2005;
Hersum et al., submitted), but as the rocks are likely to
have been preheated by intrusion of the overlying
Peneplain sill prior to emplacement of the Basement Sill,
we used a uniform initial country rock temperature
of 3008C in the model. Pressure at the top of the sill
is 100 MPa, based on an intrusion depth of 1km and
assuming a lithostatic pressure gradient of 100 MPa/km.
The bulk composition of the magma (i.e. silicate
liquid þ crystals) is composed of 60% chilled marginal
composition taken from Gunn (1966) at the temperature
at which this liquid is saturated with orthopyroxene
(11308C) plus 40% orthopyroxene of mg-number 85.
To this is added 06 wt % H2O, 001 wt % CO2 and
003 wt % S. The initial volatile contents of the Ferrar
magmas are not known, and we have used a value that
falls within the high end of water contents observed in
mafic magmas [see summary by Boudreau et al. (1997)].
This value was used to prevent free energy minimization
problems at low liquid fractions. Also, the model calculation does not go to 100% crystallization so that the actual
quantity of vapor generated is relatively small. After 90%
crystallization the liquid has lost most of its CO2 and
is degassing nearly pure H2O (as discussed below);
this composition vapor would simply continue with additional crystallization. Finally, because the chilled margin
composition of Gunn is not in equilibrium with the added
orthopyroxene and volatiles, the initial calculated solid
assemblage is actually a mix of olivine (10%) and orthopyroxene (25%). Over the course of the run, olivine is
replaced with orthopyroxene by peritectic reactions with
the silicate liquid.
Chlorine and F are treated as trace elements in the
calculation, and are assumed to have vapor^silicate
liquid partition coefficients of 100 and 01, respectively.
These partition coefficient values are within the range
observed in both natural and experimental systems across
a wide range of silicate liquid compositions (Carroll &
Webster, 1994; Saal et al., 2002). Chlorine and F are otherwise assumed to not partition into any solid phases.
Minerals that crystallize over the course of the reaction
include orthopyroxene, olivine, plagioclase, magnetite,
alkali feldspar and an immiscible Fe^S liquid. In addition
to crystallization, the mush column is allowed to compact
following the model of Shirley (1986). Any vapor that is
generated is assumed to move upward at a fixed velocity
equal to 156 cm/day; this vapor ascent velocity is one-half
the characteristic compaction velocity at the start of the
simulation. Because each time step is calculated from
the compaction velocity profile (which slows down
over time as the sill solidifies), this value was selected so
that the vapor moves about one space step for each time
step over the course of crystallization. This number is
otherwise arbitrary, as there are no models that express
how bubbles move through a crystallizing and compacting
mush column. Nor does vapor degassing or migration
affect compaction rates in the model, except for the effect
of dissolved H2O in the silicate liquid on viscosity
of the latter.
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NUMBER 7
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Fig. 7. Results of IRIDIUM model of a solidifying and degassing 350 m thick orthopyroxene-phyric sill at 11308C instantaneously intruded into
country rock initially at 3008C. Profiles are shown at time of 0, 31, 107, 210 and 432 years. From left to right, plots are of: (a) temperature in the
sill (white) and country rock (grey); (b) per cent solid; (c) mass of vapor exsolved during crystallization (wt %); (d) mole percentage of the
vapor components H2O (continuous lines), CO2 (dashed lines) and H2S (dotted line, shown at 432 years only); (e) Cl (continuous lines) and F
(dashed lines) concentrations in liquid, normalized to original concentrations. (See text for additional discussion.)
None of the results presented below are qualitatively
affected by modest changes in any of the above initial
model parameters, as long as the mush column is not
initially vapor-saturated. For example, changing the
volatile concentrations of the initial silicate liquid or
vapor velocity can change the point at which the silicate
liquid becomes vapor saturated or the time scale at which
vapor fronts move, but the relative change in CO2/H2O
exsolved over time will be similar to that discussed below.
In this respect, the main conclusions are robust.
Model results for times of 0, 31, 107, 210 and 432 years
are shown in Fig. 7. Heat loss and crystallization lead to
crystallization-induced vapor saturation (i.e. second boiling) initially occurring in the upper and lower margins as
the solidification fronts propagate in toward the center of
the sill. The extents of these zones at the labeled times
are shown in the ‘Mass vapor’ plot in Fig. 7. Although
the initial silicate liquid is relatively H2O-rich, the initial
vapor separating from both the upper and lower zones
is CO2-rich (i.e. high initial CO2/H2O ratio) owing to
the much lower solubility of CO2 (Eggler et al., 1974;
Mattey, 1991; Lowenstern, 2000). As degassing progresses
the vapor separating at any stratigraphic level becomes
more H2O-rich and the CO2/H2O ratio continually
decreases. The concentration of H2S in the vapor remains
low at all times. Spatially, this is expressed by the vapor
in the upper and lower margins being more H2O-rich
than the more interior gases at any given time, at least
during the early stages of degassing. At some time between
61 and 107 years the upper and lower vapor saturation
zones merge and the entire sill becomes vapor-saturated.
Over time, CO2 is effectively flushed from the sill such
that by 432 years the fluid is essentially all H2O with
a few mol% H2S, except in the center of the sill where
the concentration of CO2 remains high.
In the upper zone, progressive devolatilization and the
separation and loss of vapor from the top of the intrusion
to the overlying country rock results in the preferential loss
of Cl from the interstitial silicate liquid. Also, although Cl
is continually being advected through the sill in vapor
evolved and ascending from silicate melt in the lower
margin, the Cl content of interstitial liquid in the
upper zone remains approximately constant. In contrast,
the concentration of F in the interstitial liquid increases
as crystallization proceeds. The net result is a decrease
over time in the Cl/F mass ratio in the upper part of
the sill.
The generation and migration of vapor is more complicated in the lower zone. Vapor exsolved near the lower
margin ascends into hotter interstitial liquids that are not
yet vapor-saturated. This results in the resorption of vapor
in the hotter, interstitial silicate liquid. This process has
four effects. First, as Cl is preferentially partitioned into
the aqueous fluid relative to F, the resorption of vapor
with a high Cl/F mass ratio into the interstitial silicate
liquid leads to the formation of a peak in the Cl concentration in the interstitial silicate liquid that migrates upward
through the magma with time. Second, the enrichment of
volatile components in stratigraphically higher interstitial
silicate liquid owing to resorption causes this interstitial
silicate liquid to achieve volatile saturation and exsolve
a vapor phase earlier than would have occurred based
on the original volatile concentrations in this silicate
liquid fraction. This results in the lower vapor-saturated
zone moving into the center of the sill faster than
the upper zone. Third, this upward advancing ‘wet front’
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BOUDREAU & SIMON
DEGASSING IN BASEMENT SILL, ANTARCTICA
Fig. 8. Calculated profiles (from left to right) of: (a) phase mode (wt %); (b) bulk-rock MgO (wt %); (c) mass vapor (wt %); (d) the solid and
liquid velocities (cm/day); (e) the deformation of the solid matrix. All are shown at 210 years of model calculation. (See text for additional
discussion.)
produces a peak in the mass of (H2O-rich) vapor evolved
that also moves upward a short distance behind the peak
in Cl concentration in interstitial liquid. Fourth, progressive degassing from the lower margin produces more
extreme depletions in CO2 in the fluid than occurs in the
upper zone. This difference occurs because, unlike the
upper margin, the lower margin does not have a source of
CO2 from degassing interstitial liquids below it.
The migration of both the ‘Mass vapor’ and Cl peaks
continues once the upper and lower vapor saturation
zones merge. However, the peak Cl/F ratio of the interstitial silicate liquid declines over time. Although not shown,
once the Cl peak passes above the center of the sill it
encounters progressively colder magma and, as pointed
out above, this will not favor increasing the Cl concentration of apatite. Furthermore, this cooler magma contains
more evolved interstitial silicate liquid, which contains
a higher concentration of F, and hence again Cl/F ratios
are not strongly elevated.
Additional compaction effects
Shown in Fig. 8 are snapshots at 210 years of the phase
mode profile (wt%), bulk MgO, mass quantity of vapor,
solid and liquid velocities and the per cent deformation of
the solid matrix. The profiles in most of these plots show
the effects of compaction. At 210 years, the model solid
and liquid velocities in the margins are near zero.
Thus, active compaction of the mush column at this time
is slowing considerably and is occurring mainly in the
central part of the sill. The bulk MgO concentration
has evolved from a straight profile to an ‘s’-shaped profile
owing to compaction. It should be noted that, in the model
presented here, the sill was assumed to be uniform
in its initial phenocryst content. Although this is not
entirely consistent with the conclusion of Marsh (2004),
we highlight that the new model results would be superimposed on any more complicated initial distribution of phenocrysts. In the lower part of the intrusion, compaction
leads to an increase in MgO as solids displace interstitial
liquid. In contrast, for reasons discussed in some detail by
Boudreau & Philpotts (2002), in the upper part of the
sill the solid assemblage is undergoing extensional deformation. Briefly, this occurs because the increasing solidification of interstitial silicate liquid towards the upper
contact decreases the permeability and hence the solid
velocity decreases upward in the upper half of the sill.
In effect, solids in the center of the sill can move down
faster than those near the upper margin. The effect is
a zone of extension that is maximized at about two-thirds
up in the sill. Although it is beyond the scope of this study,
we note that in the Dais intrusion the tensional region
is characterized by the appearance of pronounced
mafic-felsic layering.
Finally, it is noted that the migration of exsolved vapor
affects the compaction-driven silicate liquid velocity.
In general, the silicate liquid would have a positive
velocity value as the liquid moves upward in response
to solids moving downward (i.e. compacting). However,
vapor ascends faster than silicate liquid, hence there is
a local negative liquid velocity as a ‘backflow’ of silicate
liquid displaces ‘lost’ vapor. This is most evident where
a negative liquid velocity peak is associated with the
vapor mass peak in the lower part of the sill, but it is
also present in the upper part of the sill.
CONC LUSIONS
Apatite compositions and numerical modeling both
suggest that degassing in a mafic sill can be relatively
complicated. Like the magma itself, the vapor evolved
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JOURNAL OF PETROLOGY
VOLUME 48
from a cooling magma can undergo a protracted period of
compositional evolution. This affects the estimation
of volatile abundances of elements such as S, Cl and F.
Degassing from the lower margin of the sill can influence
strongly the volatile-element budget of the overlying
magma column as crystallization proceeds. Variation in
halogen ratios of apatite through the vertical extent of
a sill provide a convenient, if difficult, method for elucidating its volatilization history. The model results agree
broadly with the fluid inclusion record preserved in the
Stillwater Complex (Hanley et al., 2005).
AC K N O W L E D G E M E N T S
The senior author would like to thank Bruce Marsh for
his efforts to make the Field Laboratory Workshop in the
McMurdo Dry Valleys, Antarctica, a great success and
for the chance to participate in the study of these very
interesting rocks. This paper was significantly improved
by comments from W. Bohrson, D. Geist, J. Hanley and
an anonymous reviewer. The work was supported by
National Science Foundation Grant EAR 04-07928.
S U P P L E M E N TA RY DATA
Supplementary data for this paper are available at
Journal of Petrology online.
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APPENDIX
The modeling described in this paper uses an updated
version of the IRIDIUM program (Boudreau, 2003).
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DEGASSING IN BASEMENT SILL, ANTARCTICA
Fig. A1. Composition of vapor phase (in mole fractions) as a function of log f(O2). Initial bulk composition is MORB liquid with 20
wt % H2O, 015 wt % CO2 and 01 wt % S, equilibrated at 12258C,
1000 bars. Log f(O2) varies between QFM þ 3 and QFM ^ 3 (where
QFM is the quartz^fayalite^magnetite buffer).
Fig. A3. Calculated sulfur concentration at sulfide saturation in
liquids of basaltic composition as a function of FeO total concentration at 12008C, 1bar and f(O2) ¼ Ni^NiO. Samples taken from
Haughton et al. (1974).
Fig. A2. Volatile concentrations in (a) vapor and (b) silicate liquid
for fractional degassing. Initial MORB liquid with 20 wt % H2O,
015 wt % CO2 and 01 wt % S at 12258C. Liquid undergoes degassing
with loss of vapor while undergoing an isothermal pressure change
from 2001 to 1bar.
Fig. A4. Sulfur concentration in silicate liquid at sulfide saturation as
a function of (a) log f(O2) and (b) log f(S2) at 12008C and 1bar.
Sample 8F1 from Haughton et al. (1974); FeOtotal is 262 wt %.
This program uses a simplified version of the MELTS
program (Ghiorso et al., 1994; Ghiorso & Sack, 1995) for
phase equilibrium determinations that are coupled with
mass and heat transfer models. The main simplification
from MELTS is the use of simpler solid solution models
for some of the minerals (Boudreau, 1999). The updates
described here include the addition of CO2 and FeS as
system components, which allows the model to exsolve
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both a C^O^H^S fluid and to precipitate sulfides and
C-bearing phases from the liquid.
CO2 solubility in IRIDIUM essentially incorporates the
model of Papale (1999). It has been modified as follows: the
CO2^oxide liquid interaction parameters have been recalibrated to the liquid components and interaction parameters used in MELTS (the Papale model used those of the
program SILMIN, an earlier version of MELTS). The thermodynamic properties of pure CO2 and all other gas
species are calculated by the general equation of state for
supercritical gases of Duan et al. (1996). Gases are otherwise
assumed to mix ideally. Possible gas species include H2O,
CO2, CO, CH4, H2S, and SO2; however, for the modeling
done here only H2O, CO2 and H2S are considered. Graphite
(C), calcite and dolomite are possible precipitating solids.
Finally, because of its low solubility at low to moderate
pressures, the model treats the CO2 liquid component as
not having any effect on the other liquid species other than
acting as a dilutant to the mole fraction, X.
The program uses liquid FeS (liquid troilite) for the
silicate liquid component for sulfur. Sulfur fugacity is calculated by the method of Wallace & Carmichael (1992).
Immiscible sulfide liquid is assumed to be an ideal solution
of two components, FeS and FeO. Other possible sulfide
phases that can precipitate are pyrrhotite, pyrite and
troilite.
FeS^oxide liquid interaction parameters were calibrated
(with some arbitrary adjustment because of uncertain
sulfide and silicate liquid compositions in the experimental data) from the data of Haughton et al. (1974),
Wendlandt (1982), Mavrogenes & O’Neill (1999) and
O’Neill & Mavrogenes (2002). Because these studies invariably tabulated total iron only, FeO and Fe2O3 are calculated by the method of Kress & Carmichael (1991) based
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JULY 2007
on experimental f(O2), temperature and bulk composition.
These experimental studies are mostly of basalt^
intermediate composition, but the lower solubility calculated for more silicic compositions appears reasonable.
Above 1 GPa, sulfide saturation values are significantly
lower than expected. As for CO2, the interaction parameters do not affect the calculation of the chemical
potential of other liquid components other than a
minor effect on mole fractions.
Examples of program output are shown in Figs A1^A4.
Figure A1 shows an example of gas speciation as a function
of f(O2) for a gas-saturated mid-ocean ridge basalt
(MORB) magma. It is necessary to note an important
caveat regarding oxygen and sulfur fugacity as calculated
by MELTS/IRIDIUM; that is, the liquid interaction
parameters used by MELTS were not explicity
calibrated against oxygen fugacity. Because of this,
when f(O2) is calculated from reactions between liquid
components used by MELTS, it is not the same as that
calculated by the empirical Kress & Carmichael (1991)
model incorporated into MELTS. Nonetheless, gas speciation for S-bearing species is broadly consistent with the
model results of Shi (1992), for example.
Figure A2 shows the changes in both gas and magma composition as a vapor-saturated MORB magma undergoes
polybaric decompressional fractional degassing from 2001
to 1bar. This can be compared with similar plots of Papale
(1999).
Figure A3 shows the S concentration at sulfide saturation
(SCSS) as a function of total iron concentration in the
liquid for a selection of liquids taken from Haughton et al.
(1974), and Fig. A4 shows the SCSS as a function of sulfur
and oxygen fugacity, again for a sample from Haughton
et al. (1974).
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