JOURNAL OF PETROLOGY
VOLUME 53
NUMBER 9
PAGES 1837^1866
2012
doi:10.1093/petrology/egs035
High-temperature, low-H2O Silicic Magmas of
the Yellowstone Hotspot: an Experimental Study
of Rhyolite from the Bruneau^Jarbidge Eruptive
Center, Central Snake River Plain, USA
RENAT R. ALMEEV1*, TORSTEN BOLTE1, BARBARA P. NASH2,
FRANCOIS HOLTZ1, MARTIN ERDMANN1 AND
HENRIETTA E. CATHEY2
1
INSTITUTE OF MINERALOGY, LEIBNIZ UNIVERSITY OF HANNOVER, CALLINSTRAE 3, 30167 HANNOVER, GERMANY
2
DEPARTMENT OF GEOLOGY AND GEOPHYSICS, UNIVERSITY OF UTAH, SALT LAKE CITY, UT 84112-011, USA
RECEIVED OCTOBER 7, 2011; ACCEPTED MAY 3, 2012
ADVANCE ACCESS PUBLICATION JULY 4, 2012
The phase relations have been investigated experimentally at 200
and 500 MPa as a function of water activity for one of the least
evolved (Indian Batt Rhyolite) and of a more evolved rhyolite
composition (Cougar Point Tuff XV) from the 12·8^8·1 Ma
Bruneau^Jarbidge eruptive center of the Yellowstone hotspot.
Particular priority was given to accurate determination of the water
content of the quenched glasses using infrared spectroscopic techniques. Comparison of the composition of natural and experimentally
synthesized phases confirms that high temperatures (49008C) and
extremely low melt water contents (51·5 wt % H2O) are required
to reproduce the natural mineral assemblages. In melts containing
0·5^1·5 wt % H2O, the liquidus phase is clinopyroxene
(excluding Fe^Ti oxides, which are strongly dependent on fO2),
and the liquidus temperature of the more evolved Cougar Point Tuff
sample (BJR; 940^10008C) is at least 308C lower than that of
the Indian Batt Rhyolite lava sample (IBR2; 970^10308C). For
the composition BJR, the comparison of the compositions of the natural and experimental glasses indicates a pre-eruptive temperature
of at least 9008C. The composition of clinopyroxene and pigeonite
pairs can be reproduced only for water contents below 1·5 wt %
H2O at 9008C, or lower water contents if the temperature is higher.
For the composition IBR2, a minimum temperature of 9208C is
necessary to reproduce the main phases at 200 and 500 MPa. At
200 MPa, the pre-eruptive water content of the melt is constrained
*Corresponding author.
E-mail: [email protected]
in the range 0·7^1·3 wt % at 9508C and 0·3^1·0 wt % at 10008C.
At 500 MPa, the pre-eruptive temperatures are slightly higher
(by 30^508C) for the same ranges of water concentration.
The experimental results are used to explore possible proxies to constrain the depth of magma storage. The crystallization sequence of
tectosilicates is strongly dependent on pressure between 200 and
500 MPa. In addition, the normative Qtz^Ab^Or contents of
glasses quenched from melts coexisting with quartz, sanidine and
plagioclase depend on pressure and melt water content, assuming
that the normative Qtz and Ab/Or content of such melts is mainly
dependent on pressure and water activity, respectively. The combination of results from the phase equilibria and from the composition
of glasses indicates that the depth of magma storage for the IBR2
and BJR compositions may be in the range 300^400 MPa
(13 km) and 200^300 MPa (10 km), respectively.
KEY WORDS: Yellowstone hotspot; rhyolite; phase equilibria; crystallization experiments; Bruneau^Jarbidge; Snake River Plain
I N T RO D U C T I O N
The activity of mantle plumes has a major influence on
large-scale geodynamic processes, including the breakup
ß The Author 2012. Published by Oxford University Press. All
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JOURNAL OF PETROLOGY
VOLUME 53
and rifting of continents and attendant eruption of large
volumes of silicic magma generated by interaction of crustal materials with mantle-derived magmas. In contrast to
the numerous investigations of hotspot volcanism occurring in the oceanic lithosphere, much less attention has
been paid to hotspot volcanism within the continental
lithosphere. This may be related to the comparatively limited exposure but also to the complex interactions between
mantle- and crustal-derived melts that characterize continental hotspot volcanism. One example allowing us to gain
insights into the complex interrelationship between
mantle and crustal processes is the Yellowstone hotspot, because its activity has not been overprinted by later major
geodynamic processes. The hotspot first manifested itself
at 17 Ma, producing the Columbia River and Steens
Basalts, and the associated eruptive centers then migrated
eastwards to their current position below the Yellowstone
Volcanic Plateau (Pierce & Morgan, 1992; Perkins &
Nash, 2002; Camp & Ross, 2004). Fundamental questions
on the existence of a deep mantle plume or alternatively
of a thermal anomaly below Yellowstone, as well as on the
formation of the basalts, are still subjects of debate
(Christiansen et al., 2002; Shervais et al., 2005, 2006;
Leeman et al., 2008). However, recent tomographic studies
have imaged a major low-velocity anomaly that extends to
depths of at least 1200 km in the mantle below the
Yellowstone Volcanic Plateau (Smith et al., 2009; Obrebski
et al., 2010). Nonetheless, the formation conditions and the
source of the considerable volumes of rhyolitic magmas
that have been erupted in the last 15 Myr remain
enigmatic.
The rhyolitic volcanism related to the Yellowstone
Volcanic Province has been investigated in numerous studies including the initial descriptions of the voluminous
volcanism of the central Snake River Plain by Bonnichsen
(1982a, 1982b) and Bonnichsen & Citron (1982). Since that
time there have been many studies on the petrology of the
Yellowstone Volcanic Province, too numerous to mention
here, that are referred to in more recent publications to
which the reader is directed (Cathey & Nash, 2004, 2009;
Morgan & McIntosh, 2005; Nash et al., 2006; Bindeman et
al., 2007; Bonnichsen et al., 2008; Leeman et al., 2008;
Girard & Stix, 2009, 2010; McCurry & Rodgers, 2009;
Vazquez et al., 2009; Ellis et al., 2010; Watts et al., 2011).
Although Pb isotopic data are consistent with an ancient
source for the Yellowstone Volcanic Province rhyolites, Sr
and Nd isotopic data preclude direct melting of the
Precambrian basement that underlies much of the volcanic
province (Nash et al., 2006; Leeman et al., 2008). Yellowstone Volcanic Province rhyolites contain a significant juvenile componentçeither younger crustal source material
(such as the Idaho Batholith) or gabbro crystallized from
the basalts of the Snake River Plain or both. Similar to
the basalts, the isotopic compositions of the rhyolites
NUMBER 9
SEPTEMBER 2012
indicate a change in the lithospheric composition between
the eastern and western Snake River Plain, which is
marked by the transition from accreted oceanic terranes
(in the west) to Precambrian basement (in the east). Rhyolite compositions range from low to high silica (68^76 wt
%), medium to high K2O, and a distinctive feature is the
anhydrous mineral assemblage, consistent with high magmatic temperatures. Based on classical geothermometers
(compositions of mineral pairs, zircon saturation), a general decrease of temperature with time has been observed
for the Snake River Plain rhyolites. The earliest erupted silicic magmas in the western part of the Snake River Plain
(16^17 Ma) were probably extremely hot (up to 10808C;
Perkins & Nash, 2002). Temperatures in the central Snake
River Plain, including the Bruneau^Jarbidge and Twin
Falls eruptive centers, were typically in the range
900^10008C (Honjo et al., 1992; Cathey & Nash, 2004,
2009; Ellis et al., 2010), whereas younger rhyolites in the
eastern part (particularly Heise and Yellowstone) are
more evolved, have lower pre-eruptive temperatures (typically 800^9008C) and at Yellowstone some contain biotite
or hornblende (Christiansen, 2001; Watts et al., 2011), indicating a change of magma storage conditions (or water
content) in both time and space. For example, considering
that several rhyolitic suites contain plagioclase, sanidine
and quartz, the water contents of the early, high-temperature rhyolitic magmas are expected to be very low
(e.g. Holtz et al., 2001) and should be significantly higher
in the younger, low-temperature magmas, particularly
those including hornblende in the phase assemblage.
Although many workers have attributed the anhydrous
mineral assemblage in Snake River Plain rhyolites to relatively low magmatic water contents, virtually nothing is
known about water concentrations in the pre-eruptive
magma reservoirs and how those concentrations may
have varied with time (Ellis et al., 2010).
Additionally, there are nearly no data on the depth of
the main rhyolitic magma reservoirs. The experimental
determination of phase equilibria at high pressure and
temperature is one of the methods commonly applied to
constrain pre-eruptive conditions (e.g. Rutherford et al.,
1985; Scaillet & Evans, 1999; Holtz et al., 2005). This study
reports the first attempt to constrain experimentally the
magma storage conditions of the high-temperature rhyolitic magmas of the Bruneau^Jarbidge eruptive center in
the central Snake River Plain.
B RU N E AU ^ J A R B I D G E
VO L C A N I C C E N T E R
The Bruneau^Jarbidge volcanic center derives its name
from the Bruneau and Jarbidge Rivers that flow northward
across the south^central Snake River Plain. Their deeply
incised canyons provide outstanding vertical exposures of
1838
ALMEEV et al.
SILICIC MAGMAS OF YELLOWSTONE HOTSPOT
up to 475 m of large-volume, silicic ignimbrites and rhyolite
lava flows. The eruptive activity of the center can be
divided into two episodes. The first episode, from 12·8 to
10·5 Ma, was dominated by a sequence of 10 largevolume, ash-flow eruptions, several in the supereruption
category with volumes in excess of 500^1000 km3 (Miller
& Wark, 2008). These 10 eruptive units constitute the
Cougar Point Tuff (CPT) as initially described by
Bonnichsen (1982a) and Bonnichsen & Citron (1982). The
explosive phase at Bruneau^Jarbidge was followed by
rhyolite lava flows, some with volumes of 75 to 200 km3,
erupted over the interval from 11 to 8 Ma (Bonnichsen,
1982b). Although a few lava flows are intercalated with
the youngest of the major ash-flow units, the later activity
at Bruneau^Jarbidge was largely effusive. The total
volume of erupted material is estimated to be between
7000 and 10 000 km3, not including caldera infill and
distal ash-fall tuffs (Perkins & Nash, 2002; Bonnichsen
et al., 2008).
The Bruneau^Jarbidge system exhibits a systematic temporal variation in the chemical composition of the silicic
magmas in which the initial eruptions are more silicic,
and the magmas become less evolved (more mafic)
with time (Bonnichsen et al., 2008). Both ash flows and
lavas are characterized by high eruption temperatures of
900^10008C, and these high temperatures persist through
the waning stages of lava-dominated volcanism (Cathey
& Nash, 2004, 2009). In addition to elevated temperatures
and large volumes, other unique features of the central
Snake River Plain volcanism, including the absence of
abundant pumice, rheomorphism, intense welding and
fusion of underlying ash-fall layers, among others, led
Branney et al. (2008) to define a new category of volcanism
termed ‘Snake River (SR)-type volcanism’. Also, although
not a focus of this study, it is important to note that the
Bruneau^Jarbidge eruptive center constitutes the largest
low d18O silicic volcanic province known on Earth in
which all of the eruptive products show significant 18O depletion, with an average d18O in zircon from the Cougar
Point Tuff of 1· 0% (Cathey et al., 2011).
EXPERIMENTS
Starting material
The natural samples of the Bruneau^Jarbidge eruptive
center are predominantly composed of glass (largely devitrified, except in basal and upper vitrophyre zones), and
phenocryst contents typically are in the range of 5^15%
in ash-flow tuffs and 10^20% in rhyolite lavas. The phenocryst assemblage consists of plagioclase (Plag), sanidine
(San), quartz (Qtz), augite (Aug), pigeonite (Pig), magnetite (Mt), ilmenite (Ilm), and fayalite (Fa). However, many
samples contain only a subset of these minerals (Cathey
& Nash, 2004). Accessory minerals that have been identified are apatite and zircon.
Two samples with slightly different compositions
(Table 1) were used for the experimental investigation. The
most evolved sample of the two (BJR) is from the basal
vitrophyre of member XVof the Cougar Point Tuff exposed
in the Jarbidge River drainage (Bonnichsen, 1982a;
Bonnichsen & Citron, 1982), also referred to as XVj by
Perkins & Nash (2002) to distinguish it from the compositionally different XVb in the Bruneau River drainage,
and as unit 9j (Cathey & Nash, 2004). XVj and XVb are
the last of the 10 major ash-flow eruptions from the
Bruneau^Jarbidge eruptive center, and have an approximate age of 10·5 Ma based on tephrochronology of fallout
tuffs (Perkins et al., 1995; Perkins & Nash, 2002). Chemical
and mineralogical data on this unit of the Cougar Point
Tuff have been provided by Cathey & Nash (2004), including a temperature estimate for pyroxene equilibration of
9368C. The second sample, IBR2 (Indian Batt Rhyolite), is
one of the least evolved effusive eruptive units of the
Bruneau^Jarbidge eruptive center (Bonnichsen, 1982b). It
underlies one of the youngest of the rhyolite lavas, the
Dorsey Creek rhyolite (8·2 0·2 Ma), and has an age of
9·3 0·7 Ma based on in situ analysis of U^Th^Pb in
zircon by SIMS (Nash B.P., unpublished data). Whole-rock
compositions for BJR and IBR2 are given inTable 1.
The basal vitrophyre sample, BJR, is composed of a dark
glassy matrix and phenocrysts of plagioclase (4^5 vol.
%), augite, pigeonite and Fe^Ti oxides, with minor
amounts of sanidine, fayalite and quartz (Fig. 1). The total
phenocryst proportion is estimated to be 10 vol. %.
To facilitate direct comparison with the experimental
run products, phenocrysts in the samples used for the
experiments were analyzed using the electron microprobe
facility at Hannover (see below). In BJR, plagioclase
composition is in the range An26^34 with Or in the range
Or7^11. The composition of sanidine varies from
An2Or60Ab38 to An4Or49Ab47. Augite compositions vary
from Fs48Wo35En17 to Fs50Wo34En16. Pigeonite compositions vary from Fs69Wo13En18 to Fs66Wo10En24. Olivine is
close to fayalite with Fo12Fa88.
The less evolved rhyolite lava sample, IBR2, contains
plagioclase, augite, and pigeonite as phenocryst phases
and lacks quartz and sanidine. The composition of the
pyroxenes is more Mg-rich than those of BJR. Augite compositions vary from Fs28Wo40En32 to Fs33Wo36En31.
Pigeonite compositions vary from Fs51Wo9En40 to
Fs53Wo8En39. Plagioclase composition varies from An33 to
An40 (with Or in the range Or3^8). In their study of pyroxenes in rhyolite lavas of the Bruneau^Jarbidge center,
Cathey & Nash (2009) reported a crystallization temperature of 975 118C for the Indian Batt rhyolite based on
pyroxene thermometry.
Natural plagioclase compositions in the products of the
Bruneau^Jarbidge eruptive center show a total range of
An24^36 Or6^10 Ab57^67 for the CPT (Table 2, Cathey &
1839
JOURNAL OF PETROLOGY
VOLUME 53
Table 1: Chemical compositions of the studied rhyolites
(normalized to 100%)
Indian Batt Rhyolite
Sample:
n:
SiO2
Cougar Point Tuff
IB THT-02
IBR2
IBR2
BJR
CPTXVj
BJR
rock*
starting
gm
starting
9j rock*
gm
glass
glass
glass
200
14
30
75·40
73·74
2
70·12
70·77
glass
4
74·08
9
75·45
TiO2
0·67
0·65
0·39
0·46
0·43
0·28
Al2O3
14·13
13·22
12·23
11·96
12·51
12·22
FeO
4·14
3·98
1·89
4·00
3·15
2·51
MnO
0·07
0·06
0·03
0·01
0·04
0·04
0·06
MgO
0·67
0·62
0·13
0·24
0·35
CaO
2·44
2·08
0·86
1·13
1·08
0·63
BaO
0·13
0·14
0·12
0·26
n.a.
0·19
Na2O
2·71
2·96
2·76
2·22
2·40
2·94
K2O
4·91
5·41
6·15
5·93
5·89
5·64
P2O5
0·14
0·11
0·03
0·05
0·06
0·03
*From Bonnichsen et al. (2008).
gm, groundmass; n.a., not analyzed.
Nash, 2004) and a similar range in post-CPT lavas that extends to An contents of 45^50 (mol %). The spectrum of
compositions as well as the average plagioclase composition (e.g. An content) shifts accordingly with pre-eruptive
magma temperatures determined from pyroxene thermometry and glass composition, suggesting equilibrium
with coexisting liquids. There is little evidence for xenocrysts, although antecrystic plagioclase cores that are overgrown by rims in equilibrium with coexisting liquids are
suspected in several units. Plagioclase textures are variable
in both lavas and tuffs, and include solitary, euhedral
crystals with few or no inclusions, as well as crystals with
sieved or boxy cellular cores overgrown by clear,
inclusion-free rims. Plagioclase also occurs commonly in
glomerocrystic clots with magnetite and pyroxene, but
compositions do not appear to vary according to texture.
Experimental methods
Crystallization experiments were performed using finely
powdered fused glass of natural rock powder of the samples
(BJR and IBR2). The starting glasses were prepared by
fusing the sample twice at 16008C and 1atm in air for at
least 3 h. The compositions of the glasses obtained after
the second fusion are given in Table 1. Replicate electron
microprobe analyses (200 analyses for both compositions) show that the glasses are homogeneous within analytical uncertainty and that Fe and Na losses during
fusion were minimal (compare the whole-rock analysis of
NUMBER 9
SEPTEMBER 2012
the natural samples with the microprobe analysis of the
dry glass; Table 1).
The glasses synthesized at 1atm were crushed in an
agate mortar to a grain size 5200 mm and were used as
the starting material for the crystallization experiments.
The water activity, aH2O, of the experimental charges
was varied by adding a fluid composed of a mixture of
H2O and CO2. The starting materials were sealed in gold
capsules (15 mm length, 2·8 mm internal diameter and
0·2 mm wall thickness). First distilled H2O, then CO2,
added as silver oxalate for H2O-undersaturated runs, and
finally glass were loaded into the capsules. Most runs were
fluid-saturated with a fluid/glass ratio (H2O þ CO2)/
(H2O þ CO2 þ silicate) of less than 0·11 (by weight).
However, some experiments were conducted at nominally
dry conditions (no fluid added). It is emphasized that
such experiments are not strictly water-free because (1) it
is nearly impossible to avoid adsorbed water on the surface
of the glass grains, and (2) hydrogen that can be present
in the pressure medium (gas) may diffuse through the
noble metal capsules. Thus in nominally dry experiments,
a fluid phase was not present, but the silicate melts contained small amounts of water mainly dissolved as OH
groups (0·3^1·0 wt % depending on pressure and extent
of crystallization).
Crystallization experiments were performed at 200 and
500 MPa in an internally heated pressure vessel (IHPV)
at temperatures ranging from 825 to 10258C; the experimental conditions are summarized in Tables 2 and 3. The
capsules were first pressurized and then heated up to the
run temperature with a heating rate of 508C min1. Most
experiments were conducted at low water activities and
high temperature. These conditions are consistent with
the anhydrous phenocryst assemblages and with the
geothermometry results. The run duration varied with
temperature: more than 500 h (20 days) for runs at 800
and 8508C, and more than 160 h (6 days) for runs at
900, 950 and 10008C (Tables 2 and 3).
Most experiments were conducted at intrinsic oxygen
fugacity conditions (marked as i in column ‘fH2’ in
Tables 2 and 3) using pure Ar as a pressure medium.
Under H2O-saturated conditions in this set of experiments
the oxygen fugacity was determined to be 3·3 log units
above the oxygen fugacity of the quartz^fayalite^magnetite (QFM) solid oxygen buffer (hereafter labeled
QFM þ 3·3). For comparison, analyses of magnetite^ilmenite pairs from the Cougar Point Tuff (CPT) indicate
that the natural magmas crystallized under conditions
very close to the QFM buffer (Cathey & Nash, 2004).
It should be noted that with the IHPV technique the
oxygen fugacity in the sample container decreases with
decreasing water activity (e.g. Scaillet et al., 1992), and is
close to the QFM buffer at nominally dry conditions
(Berndt et al., 2002; Botcharnikov et al., 2005).
1840
ALMEEV et al.
SILICIC MAGMAS OF YELLOWSTONE HOTSPOT
Fig. 1. Photomicrographs of the natural samples BJR9j (a) and IBR2 (b) used as starting material for the experiments. The phases that can
be observed are labeled. Aug, augite; Fa, fayalite; Mt, magnetite; Pig, pigeonite; Pl, plagioclase; Qtz, quartz; San, sanidine; Zr, zircon.
1841
1842
204
200
200
204
204
204
204
204
204
204
200
200
200
1000
1000
1000
975
975
975
950
950
950
950
925
925
925
925
168
168
168
240
240
240
336
336
336
336
504
504
504
504
384
384
168
168
504
504
504
504
336
624
624
624
504
Duration (h)
EMPA
MIR or NIR
1·27
0·76
0·38
1·08
0·61
0·48
1·26
0·69
0·39
0·29
1·73
1·01
0·80
0·58
2·68
1·05
1·10
0·92
2·22
1·78
1·56
1·12
1·14
(wt %)
(wt %)
0·07
0·42
0·46
0·59
0·23
0·00
1·34
1·17
0·21
0·81
1·71
0·03
0·59
1·46
2·05
1·09
1·30
0·99
1·27
1·37
1·16
1·25
1·65
4·83
3·56
3·32
2·20
melt H2O
melt H2O
1·27
0·76
0·38
1·08
0·61
0·48
1·26
0·69
0·39
0·29
1·73
1·01
0·80
0·58
2·68
1·05
1·10
0·92
2·22
1·78
i
i
i
i
i
i
i
i
i
i
i
i
i
i
i
i
i
i
i
i
i
i
1·12
1·56
CSPV
i
i
i
i
(bar)
fH2
1·65
4·83
3·56
3·32
2·20
for diagram
(wt %) used
melt H2O
0·13
0·05
0·01
0·10
0·03
0·02
0·12
0·04
0·01
0·01
0·21
0·08
0·06
0·03
0·43
0·14
0·10
0·07
0·31
0·22
0·18
0·10
0·19
0·89
0·61
0·56
0·30
aH2O
9·24
10·02
11·12
9·87
10·75
11·15
10·05
10·96
11·91
12·36
10·01
10·79
11·14
11·67
9·82
10·78
11·09
11·36
10·55
10·84
11·03
11·52
14·04
10·59
10·93
11
11·53
(bar)
log fO2
Gl*(99·2), Mt*(0·8)
Gl*(98), Mt*(2·1)
1·51
Gl*(98·7), Pig*(0·8), Fsp*(0·1), Mt*(0), Fa*(0·4)
Gl*(99·1), Mt*(0·9)
Gl*(99·1), Aug*(0·6), Mt*(0·4)
Gl*, Aug, Pig*, Fsp, Mt
Gl*(97·1), Aug*(0·9), Fsp*(0), Mt*(2)
Gl*(96·9), Aug*(1·2), Fsp*(0·1), Mt*(1·7)
Gl*, Aug, Pig*, Fsp*, Qtz*, Mt
Fa*(1)
Gl*(93·3), Aug*(1·2), Pig*(0·7), Fsp*(2·3), Qtz*(1·5),
Gl*(96·8), Aug*(0·8), Mt*(2·4)
Gl*(96·1), Aug*(1·5), Fsp*(0·6), Qtz*(0·9), Mt*(0·9)
0·74
0·36
1·27
0·39
0·01
1·48
0·57
0·38
0·83
1·94
1·16
Gl*, Aug*, Fsp*, Qtz*, Mt
Gl*, Aug, Pig*, Fsp*, Qtz*, Fa*
0·28
0·81
Gl*(96·5), Aug*(0·7), Mt*(2·8)
Gl*(87·4), Aug*(1·8), Fsp*(7·3), Qtz*(2·4), Mt*(1·1)
2·56
1·6
Mt*(0·4)
Gl*(67·5), Aug*(2·7), Pig*(0), Fsp*(18·6), Qtz*(10·7),
Gl*, Aug*, Pig*, Fsp, Qtz*, Mt, Fa*
1·02
1·29
Gl*(88·2), Aug*(2·4), Fsp*(6·8), Mt*(2·6)
Gl*(86·3), Aug*(3·2), Fsp*(7·4), Qtz*(1·2), Mt*(2)
Gl*(73·3), Aug*(3), Fsp*(15·9), Qtz*(5·7), Mt*(2·2)
Gl*, Aug*, Pig*, Fsp, Qtz*, Fa*
·5), Mt*(0·3), Fa*(0·2)
Gl*(78·6), Aug*(2),Pig*(1·3), Pl*(0·5) Fsp*(11·4), Qtz*(2
Gl*, Aug*, Fsp*, Qtz*, Mt*
Gl*, Aug, Fsp*, Qtz*, Mt*
Gl*(83·9), Aug*(2), Fsp*(8·7), Qtz*(2·5), Mt*(2·9)
Gl*, Aug, Pl, Fsp*, Qtz*, Mt, Fa
proportions (mass balance)
Phase assemblage and, when possible, phase
2·28
1·99
1·8
1·31
0·74
3·2
2·86
2·79
2·26
QFM
0·12
0·06
0·04
0·26
0·01
0·08
0·07
2·1
0·8
1·3
0·9
1·0
2·9
3·1
3·2
6·7
0·09
3·9
3·5
12·6
32·5
11·8
13·8
26·7
21·6
16·1
j (%)
0·08
0·11
0·21
0·11
0·10
0·41
0·51
0·43
0·03
R2
*Phases with successful microprobe measurements.
ox, experimental runs without controlled oxygen fugacity; i, IHPV intrinsic conditions (fixed hydrogen pressure); CSPV, experiments in cold seal pressure vessels
buffered by the Ni–NiO solid assemblage [see details given by Klimm et al. (2003)]. Phases: Gl, residual glass; Aug, augite; Pig, pigeonite; Bt, biotite; Amph,
amphibole; Pl, plagioclase; Fsp, K-feldspar; Qtz, quartz; Mt, magnetite; Fa, fayalite.
ox
ox
BJR-49
BJR-13
ox
BJR-57
ox
ox
BJR-54
BJR-76
ox
BJR-56
ox
ox
BJR-55
BJR-75
ox
BJR-58
ox
ox
BJR-59
ox
ox
BJR-60
200
900
900
900
900
875
875
875
875
850
825
825
825
825
T (8C)
NUMBER 9
BJR-50
ox
BJR-61
204
204
200
200
200
200
200
200
201
200
200
200
204
P (MPa)
VOLUME 53
BJR-51
ox
ox
BJR-73
ox
ox
BJR-72
BJR-23
ox
BJR-66
BJR-22
ox
CSPV
BJR-4
BJR-65
ox
BJR-69
ox
ox
BJR-70
ox
ox
BJR-71
BJR-62
ox
BJR-39
BJR-64
Conditions
Run
Table 2: Experimental conditions and phase assemblages for Cougar Point Tuff rhyolite (BJR)
JOURNAL OF PETROLOGY
SEPTEMBER 2012
ox
ox
ox
ox
red
red
red
ox
red
red
ox
ox
IBR2-13
IBR2-12
IBR2-53
IBR2-50
IBR2-51
IBR2-52r
IBR2-52
IBR2-53r
IBR2-70
IBR2-29
IBR2-30
ox
IBR2-16
ox
ox
IBR2-15
IBR2-11
ox
IBR2-18
IBR2-10
ox
IBR2-26
red
ox
IBR2-28
red
ox
IBR2-24
IBR2-35
ox
IBR2-27
IBR2-34
ox
IBR2-25
red
ox
IBR2-21
IBR2-14
ox
IBR2-20
ox
ox
IBR2-23
IBR2-17
Conditions
Run
1843
200
200
204
204
204
204
204
204
204
200
200
200
200
200
200
200
200
200
200
200
200
200
200
200
200
200
200
200
P (MPa)
1000
1000
975
975
975
975
975
975
975
950
950
950
950
930
930
930
925
925
925
925
900
900
900
900
900
875
875
825
T (8C)
168
168
288
288
240
288
288
288
240
336
336
336
336
163
163
163
504
504
504
504
168
168
168
168
168
504
504
0·56
0·50
4·84
1·70
1·00
0·77
0·73
0·54
0·45
1·19
0·67
0·57
1·88
0·86
2·42
1·75
1·03
0·82
1·69
1·75
1·36
0·35
0·75
5·24
1·49
0·64
0·04
0·06
0·33
0·21
1·50
0·62
0·14
0·98
2·65
1·32
0·49
2·11
0·87
0·88
0·21
3·51
2·80
2·42
1·30
1·07
1·78
2·94
EMPA
MIR or NIR
624
(wt %)
(wt %)
(h)
melt H2O
melt H2O
Duration
0·56
0·50
4·84
1·70
1·00
0·77
0·73
0·54
0·45
1·50
1·19
0·67
0·57
1·88
0·86
0·49
2·42
1·75
1·03
0·82
3·51
2·80
2·42
1·30
1·07
1·75
1·36
2·94
for diagram
(wt %) used
melt H2O
i
i
7·19
7·19
i
7·19
7·19
7·19
i
i
i
i
i
0·64
0·64
0·64
i
i
i
i
i
i
i
i
i
i
i
i
(bar)
fH2
0·03
0·02
0·92
0·21
0·08
0·05
0·05
0·03
0·02
0·17
0·11
0·04
0·03
0·24
0·06
0·02
0·35
0·21
0·09
0·06
0·60
0·44
0·35
0·13
0·09
0·21
0·14
0·46
aH2O
10·51
10·69
10·42
11·71
9·98
12·91
12·99
13·44
11·25
9·79
10·13
11·01
11·28
10·29
11·44
12·34
9·55
9·99
10·76
11·11
9·52
9·8
11·15
10·85
11·15
10·88
11·24
11·15
(bar)
log fO2
Table 3: Experimental conditions and phase assemblages for Indian Batt rhyolite (IBR2)
0·25
0·07
0·72
0·57
1·16
1·77
1·85
2·3
0·11
1·75
1·41
0·53
0·26
1·57
0·42
0·48
2·4
1·96
1·19
0·84
2·86
2·58
2·64
1·53
1·23
1·95
1·59
2·62
QFM
Gl*, Aug, Pig*, Pl*
Gl*(88·9), Aug*(0·2), Pig*(2·6), Pl*(8·3)
Gl*
Gl*(97·1), Aug*(2·9), Mt
0·0
11·1
0·0
2·9
16·3
15·1
22·5
8·7
14·0
20·1
12·8
10·5
19·2
4·2
16·1
5·0
9·8
14·8
16·4
j (%)
(continued)
0·04
0·00
0·59
0·28
Gl*(83·7), Aug, Pig*(5·7), Pl*(10·6), Ilm
Gl*, Aug, Pl*, Mt*
0·13
1·78
0·26
0·12
0·14
Gl*(84·9), Aug, Pig*(5), Pl*(10·1), Ilm
Gl*(77·5), Aug?, Pig*(6·2), Pl*(16·3), Ilm
Gl*, Aug*, Pig*, Pl*, Mt
Gl*(91·3), Aug*(2·1), Pl*(4·5), Mt*(2·1)
Gl*(86), Aug*(2·4), Pig*(0·1), Pl*(9·4), Mt*(2·1)
Gl*(79·9), Aug*(3·2), Pig*(0·1), Pl*(14·3), Mt*(2·5)
Gl*, Aug*, Pig*, Pl?, Fsp*, Qtz*, Mt
0·67
0·21
Gl*(89·5), Aug*(5·7), Pig*(0), Pl, Fsp*(4·8), Qtz*, Mt
1·72
0·24
0·21
0·49
0·25
Gl*(87·3), Aug*(3·2), Pl*(7·9), Mt*(1·7)
Gl*(80·8), Aug*(7·6), Pig?, Pl, Fsp*(11·6), Qtz, Mt
Gl*(95·8), Aug*(1·7), Mt*(2·5)
Gl*(83·9), Aug*(2·8), Pl*(10·7), Mt*(2·6)
Gl*, Aug*, Pl, Fsp*, Qtz, Mt?
Gl*, Aug*, Pig*, Pl*, Fsp*, Qtz, Mt*
Gl*(95), Aug*(2·4), Mt*(2·6)
Gl*(90·2), Aug*(3·2), Pl*(4), Mt*(2·7)
0·20
0·29
Gl*(85·2), Aug*(2·9), Pl*(9·1), Mt*(2·8)
R2
Gl*(83·6), Aug*(3·3), Pl*(10·7), Mt*(2·3)
Gl*, Aug*, Pl?, Fsp*, Qtz*, Mt?
Gl*, Aug*, Pl?, Fsp*, Qtz*, Mt*
Gl*, Aug*, Pl?, Fsp*, Qtz*, Mt*
Gl*, Fsp*, Qtz*, Mt*
(mass balance)
Phase assemblage and, when possible, phase proportions
ALMEEV et al.
SILICIC MAGMAS OF YELLOWSTONE HOTSPOT
1844
red
red
red
red
IBR2-58
IBR2-59
IBR2-60
IBR2-61
503
503
503
503
501
501
501
501
499
499
499
499
499
205
205
205
200
200
P (MPa)
1025
1025
1025
1025
1000
1000
1000
1000
975
975
975
975
975
1025
1025
1025
1000
1000
T (8C)
163
163
163
163
161
161
161
161
235
235
235
235
235
284
284
284
168
5·18
2·13 2·08
1·57 1·50
0·70
3·02
1·77 1·66
1·26
0·76
7·27
3·09
2·29
2·18 2·03
0·96
3·17
0·62
0·52
0·82
0·72
5·85
2·27
0·79
0·32
0·35
1·30
2·10
0·52
8·90
3·20
2·33
2·46
0·96
5·53
1·07
0·61
0·04
0·49
EMPA
MIR or NIR
168
(wt %)
(wt %)
(h)
melt H2O
melt H2O
Duration
5·18
2·08
1·50
0·70
3·02
1·66
1·26
0·76
7·27
3·09
2·29
2·03
0·96
3·17
0·62
0·52
0·82
0·72
for diagram
(wt %) used
melt H2O
32·63
32·63
32·63
32·63
53·1
53·1
53·1
53·1
53·57
53·57
53·57
53·57
53·57
18·9
18·9
18·9
i
i
(bar)
fH2
0·61
0·17
0·1
0·03
0·3
0·12
0·08
0·03
0·88
0·31
10·66
11·76
12·21
13·37
12·12
12·91
13·31
14·03
12·77
13·68
14·22
14·06
0·16
15·31
10·94
13·26
13·54
9·9
10·1
(bar)
log fO2
0·2
0·05
0·51
0·04
0·03
0·06
0·05
aH2O
0·48
1·58
2·03
3·19
1·58
2·37
2·77
3·49
1·85
2·76
3·14
3·3
4·39
0·55
2·87
3·15
0·86
0·66
QFM
Gl*
Gl*(97·3), Aug*(2·7)
Gl*(98·6), Aug*(1·4), Pig*(0)
Gl*(84·4), Aug*, Pig*(4·6), Pl*(10·7), Qtz*(0·3)
Gl*
Gl*(98·4), Aug*(1·6)
Gl*(97·9), Aug*(0·8), Pig*(1·2)
Gl*(60·2), Aug(0·9), Pig*(4·2), Fsp*(17), Qtz*(11·1), Ilm(0·4), Pl?
Gl*
Gl*(95·3), Aug*(4·7)
Gl*(96·8), Aug*(2), Pig*(1·2)
Gl*(96·1), Aug*(3·9), Ilm
Gl*(77·9), Aug, Pig*(5·6), Fsp*(14·8), Qtz*(1·7), Ilm, Pl?
0·00
0·82
0·26
0·01
0·00
0·07
0·18
0·02
0·00
2·22
0·30
0·24
0·57
0·00
2·15
Gl*(98·1), Aug*(1·9)
Gl*
0·00
0·08
0·09
R2
Gl*, Aug, Pig*, Pl*
Gl*(92·1), Aug*(0·7), Pig*(0·6), Pl*(6·1), Mt*(0·6)
Gl*(92·4), Aug*(0·4), Pig*(1·2), Pl*(5·7), Mt*(0·4)
(mass balance)
Phase assemblage and, when possible, phase proportions
0·0
2·7
1·4
15·6
0·0
1·6
2·1
39·8
0·0
4·7
3·2
3·9
22·1
0·0
1·9
0·0
7·9
7·6
j (%)
VOLUME 53
NUMBER 9
*Phases with successful microprobe measurements.
ox, experimental runs without controlled oxygen fugacity; red, experimental runs with controlled oxygen fugacity (Shaw membrane); i, IHPV intrinsic conditions
(fixed hydrogen pressure); mineral abbreviations are as in Table 2.
red
red
IBR2-69
IBR2-65
red
IBR2-68
red
red
IBR2-67
red
red
IBR2-66
IBR2-63
red
IBR2-49
IBR2-64
red
IBR2-48
red
red
IBR2-38
IBR2-62
ox
IBR2-31
red
ox
IBR2-32
IBR2-71
Conditions
Run
Table 3: Continued
JOURNAL OF PETROLOGY
SEPTEMBER 2012
ALMEEV et al.
SILICIC MAGMAS OF YELLOWSTONE HOTSPOT
Other experiments performed at reduced conditions
were conducted in an IHPV pressurized with a mixture
of Ar and H2 gases (the maximum H2 pressure given in
the IHPV before heating was 7·5 bar). Hydrogen diffuses
through the noble metal inside the capsules, and at 7008C
osmotic equilibrium is reached within 2 h using Au capsules (Schmidt et al., 1997). If water is present in the experimental charge, the oxygen fugacity is controlled by the
equilibrium reaction for water formation (H2 þ ½
O2 $ H2O). As a result, at a given fH2, the fO2 decreases
with decreasing water activity in the experimental charge.
The calculation of fO2 is based on the equation of Schwab
& Ku«stner (1981) [for further details see Botcharnikov
et al. (2005)]. The fH2 prevailing in the IHPV at high P
and T was controlled with a Shaw membrane (Berndt
et al., 2002). Various oxygen fugacities were obtained by
varying the proportions of H2 and Ar in the pressure
medium (see column ‘fH2’ in Tables 2 and 3). The fO2 in
the water-undersaturated experiments can be estimated
using the relation log fO2 ¼ log fO2 (at aH2O ¼1) þ 2log
aH2O (Scaillet & Evans, 1999; Botcharnikov et al., 2005),
where aH2O is determined from the water concentration in the melt following the model of Burnham
(1994). We estimate that the overall error in the
calculated fO2 is about 0·2 log units (Botcharnikov
et al., 2005).
Analytical methods
After completion of the experiments, the capsules were reweighed to ensure that there had been no loss of water
and CO2 at high pressure and temperature. No weight
loss was observed in the experiments listed in Tables 2 and
3. To check for the presence of excess fluid, capsules were
weighed and cooled in liquid nitrogen. After cooling, the
capsules were punctured immediately and reweighed to
check the weight loss caused by the release of CO2. The
capsules were subsequently heated in an oven (1008C) for
10 min to determine the weight loss of H2O in the capsule.
However, the amounts of fluid measured were too low to
be able to estimate the XH2O of the fluid accurately by
this method. The run products were examined by optical
microscope and the compositions of the mineral phases
and glasses were analyzed with a Cameca SX100 electron
microprobe (Hannover laboratory) using the following
analytical conditions for crystalline phases: accelerating
voltage of 15 kV, beam current of 15 nA, counting time 10 s
on peak for all elements. Glasses were analyzed using a defocused beam (5^10 mm beam diameter) with 15 kV voltage
and reduced beam current of 8 nA to reduce loss of Na
and K. The elements Na, K, Si, Ca and Fe were analyzed
first on each spectrometer. Ba and P in glasses were analyzed with 40 nA beam current to improve the detection
limit. Counting times varied from 8 to 30 s.
The exact determination of the water concentration in
the glasses is crucial to this study considering that the
natural systems are water-undersaturated and that water
activity is an important parameter influencing phase stabilities. Instead of using the classical ‘by-difference’ technique
(Devine et al., 1995) based on the analytical totals determined by electron microprobe analysis (EMPA), which
leads to relatively high uncertainties of 0·5^0·7 wt %
H2O (e.g. Parat et al., 2008), the water concentrations in
most experimental glasses were determined from
IR-absorption spectra using the Beer^Lambert law. In
some cases, usually experimental products with high crystallinity, an area devoid of crystals with a sufficient size
for IR measurements could not be found. For these experiments, the ‘by-difference’ technique was used to estimate
H2Omelt. Spectra were collected in the near-infrared
(NIR) and mid-infrared (MIR) range with a Bruker IFS
88 Fourier transform infrared (FTIR) spectrometer
coupled with an IR microscope A590. The microscope
and the sample zone were flushed with dry air. The spot
size was fixed at 80 mm 80 mm for most samples. This
relatively small aperture was necessary to select
crystal-poor zones in the experimental products. In some
cases (highly crystalline samples) the spot size was reduced
to 50 mm 50 mm to accommodate the crystal distribution.
The operating conditions for NIR were: W (tungsten)
light source, CaF2 beam splitter, MCT (HgCdTe) detector,
2 cm1 spectral resolution, spectral range 6000^0 cm1, 50
scans accumulated per measurement. The operating conditions for MIR were: glow bar light source, KBr beam splitter, MCT (HgCdTe) detector, 4 cm1 spectral resolution,
spectral range 13 000^0 cm1, 100 scans accumulated per
measurement.
Doubly polished samples of 60 mm thickness for MIR
and of 200 mm thickness for NIR spectroscopy were prepared from the experimental products. The thickness of
the samples was measured at five points using a digital micrometer with an estimated uncertainty of 0·0003 cm.
Three measurements per sample were performed to account for possible variations in thickness owing to polishing. For calculation of the water content, the density of
the glass was estimated to be constant at 2350 kg m3, considering (1) that the density variation of rhyolitic glasses
with changing H2O content at 200 MPa is minor
(Withers & Behrens, 1999) and (2) that the error in the determination of the exact density results in larger uncertainties in the water concentration than the density variation
itself. In addition, when compared with uncertainties on
the thickness determination, the possible variation of
water concentrations owing to changes in density are negligible (10 times less). Molar absorption coefficients of 1·41
and 1·66 l cm1 mol1 for the 4520 and 5230 cm1 bands,
respectively, were used to calculate the total water contents
from NIR measurements [coefficients were determined
for rhyolitic glasses by Withers & Behrens (1999)]. The
water contents for MIR measurements (3550 cm1) were
1845
JOURNAL OF PETROLOGY
VOLUME 53
calculated using a molar absorption coefficient of
78 l cm1 mol1. This value was determined from glasses
containing less than 2 wt % H2O synthesized in experiments above the liquidus in which the water content was
determined by Karl Fischer Titration [for this method see
Behrens et al. (1996)]. This value is slightly higher than
most of the molar absorption coefficients mentioned in the
literature, which range from 61 to 70 (Mandeville et al.,
2002). However, the dataset published by Mandeville et al.
(2002) does not contain data for high-silica rhyolites. In
addition, the data clearly show that the molar absorption
coefficients increase from basaltic to dacitic melts.
For every sample three to five single measurements were
performed and different locations in the sample were
chosen to check for homogeneity. A representative example
of three MIR measurements from a single sample is
shown in Fig. 2 together with NIR spectra collected for
samples with different water contents. The average values
of such measurements are used to calculate the water content of the glasses, with a standard deviation usually less
than 0·02 wt %.
Experimental results
Phase assemblage
The phase assemblages observed in the experimental products are given in Tables 2 and 3 (compositions BJR and
IBR2, respectively). The compositions of minerals that
could be analyzed by EPMA are reported in Tables 4
and 5 (compositions BJR and IBR2, respectively) in
terms of the main mineral components (for detailed analyses, see table in Supplementary File SM-1, available for
downloading at http:///www.petrology.oxfordjournals
.org). Crystalline phases identified include olivine, augite,
pigeonite, sanidine, plagioclase, quartz, magnetite and ilmenite (Tables 2 and 3). In two experiments amphibole
and biotite were also detected. In some experiments, the
size of the crystalline phases was too small to be analyzed
accurately by EMPA. Figure 3 shows the products that
were obtained at 9758C and 500 MPa and is representative
of the typical experimental products obtained in this
study. Experiments conducted with relatively high water
contents in the melt or at high temperature have low crystal contents and a homogeneous distribution of the mineral
phases throughout the glass matrix (mainly pyroxenes,
plagioclase and oxides). Experiments with higher crystallinity, often containing high proportions of quartz and
feldspar, show a heterogeneous distribution of the mineral
phases. Evidently the tectosilicate phases nucleated along
specific surfaces that correspond to the initial boundaries
between the grains of the dry starting glass powder inserted into the capsule. This type of experimental texture
was obtained in experiments conducted at low temperature or in nearly dry experiments where the difference between the experimental temperature and the liquidus
NUMBER 9
SEPTEMBER 2012
temperature is higher than in the experiments with a
homogeneous phase distribution.
At given P^Tconditions, the melt fraction decreases and
the proportion of crystalline phases increases with decreasing water content of the melt (hereafter labeled as
H2Omelt) (Fig. 3, Tables 2 and 3). As expected, the crystal
fraction also increases with decreasing temperature (at a
given H2Omelt). The phase compositions (SiO2, TiO2,
Al2O3, FeO*, MnO, MgO, CaO, Na2O, K2O, and P2O5)
were used to calculate modes for all charges, based on a
least-squares fit of the starting composition using the approach of Stormer & Nicholls (1978). Calculated modes
are given in Tables 2 and 3 (wt %), together with the sum
of the squares of the residuals to the fit (r2). Most fits
were good with r250·2. Fits were poor for some products
with high crystallinity or for experiments in which the
compositions of phases could not be analyzed accurately.
In these cases the proportions are not given in the tables.
The stability field of each mineral phase is depicted in
Figs 4^6 as a function of temperature and water content
of the melt. It is emphasized that the phase diagrams in
Figs 4^6 are not obtained at a specific oxygen buffer (experiments reported in Figs 4^6 have been conducted at intrinsic conditions), and that conditions are more reducing
at low water activities than at high water activities. For
clarity, the estimated oxygen fugacity for a given H2Omelt
is shown in Figs 4^6 (black arrows).
BJR (200 MPa)
Magnetite is the liquidus phase and the only phase present
at 10008C for H2Omelt40·5. For very low H2Omelt
(50·5 wt % H2O), fayalite, pigeonite and magnetite were
observed at 10008C and below (except for Fa, which was
not detected at 9758C); quartz is observed at 9508C for
H2Omelt of 0·4; at these very low H2O contents plagioclase was not observed at temperatures of 9508C and
above. At H2Omelt41wt %, augite and sanidine are the
first phases that crystallize after magnetite, followed by
quartz. Quartz is stable at 8258C up to nearly
water-saturated conditions. Plagioclase was detected in
two experiments at 8258C and H2Omelt 2·2 wt % and at
8508C and H2Omelt 1·65 wt %, and pigeonite was not
observed in experiments with H2Omelt41·1wt %.
Amphibole and biotite were identified in one experiment
at 8258C and at H2Omelt 4·8 wt %, but the phases were
too small to be analyzed accurately with microprobe.
IBR2 (200 MPa)
Experiments conducted at intrinsic oxygen fugacity are
shown in Fig. 5. At low H2Omelt (51·3 wt % H2O) the
phases observed at 10008C are augite, magnetite, plagioclase and pigeonite. Sanidine and quartz crystallize at
lower temperatures (75^508C below the crystallization of
plagioclase). Magnetite is not observed in the two experiments with the lowest H2Omelt at 10008C, which may be
1846
ALMEEV et al.
SILICIC MAGMAS OF YELLOWSTONE HOTSPOT
Fig. 2. MIR (mid-infrared) and NIR (near-infrared) spectra of glasses obtained in runs at 500 MPa and 9758C. The MIR spectra were taken
at three locations on an experimental product (IBR2-66) and illustrate the reproducibility of measurements in highly crystalline samples. The
run product of this experiment is shown in Fig. 3 (1 wt % H2O). The NIR spectra of run products of several samples with different H2O contents
are shown.
related to the low oxygen fugacity of these runs. Pigeonite
has not been detected at H2Omelt41·3 wt %. The crystallization sequence for a given H2Omelt41·3 wt % is
augite, followed by plagioclase, sanidine and quartz. The
curve showing the crystallization of augite (Cpx-in) is difficult to constrain, but the experiments at 9258C
(H2Omelt 2·4 wt %) and 9008C (H2Omelt 3·5 wt %)
contain very small amounts of augite, and the boundary
curve is expected to lie at slightly higher temperatures.
There are no experiments allowing us to determine exactly
the crystallization sequence of sanidine and quartz in the
high H2Omelt range. In contrast to the BJR composition,
fayalite was not observed. Biotite was detected in one experiment at 8258C and at H2Omelt 2·9 wt % but was
too small to be analyzed accurately.
Nine experiments (1025, 975, 9258C) were conducted at
conditions more reducing than those obtained at intrinsic
oxygen fugacity. The oxygen fugacity is lower by 3^0·4
log units when compared with intrinsic conditions. The results are given in Table 3 and, except for oxide phases in
five experiments, there is no difference between the phase
assemblages of experiments conducted at intrinsic fO2 conditions and experiments conducted at more reduced conditions (same assemblages at given T and H2Omelt). In five
1847
1848
Aug*
En(27·9)Wo(34)Fs(38·2)
BJR-49
BJR-50
En(27·8)Wo(6·6)Fs(65·6)
En(23·8)Wo(11·4)Fs(64·8)
An(9·6)Or(54)Ab(29·3)Cn(7·1)
Fsp*
An(7·2)Or(59·4)Ab(25)Cn(8·5)
X Usp(0·3)
Mt*
X Usp(0·3)
X Usp(0·3)
Mt*
X Usp(0·2)
BJR-13
*Detected phases without microprobe measurements and phase calculations.
X Usp(0·4)
BJR-76
Fo(18·6)Fa(79·4)
Qtz(34·1)Or(36)Ab(19·4)An(4·7)
Qtz(32·6)Or(35·9)Ab(19·1)An(4·3)
Qtz(32·4)Or(36)Ab(19·6)An(4·1)
Qtz(32·9)Or(35·7)Ab(18·3)An(4·9)
Qtz(31·8)Or(35·9)Ab(20·5)An(3·1)
Qtz(33·8)Or(35·3)Ab(20)An(3·8)
Qtz(33·7)Or(35·7)Ab(19·9)An(4·4)
Qtz(34)Or(36)Ab(19·1)An(4·4)
Qtz(33·2)Or(35·8)Ab(19·5)An(4·1)
Qtz(32·7)Or(36·5)Ab(19·1)An(4·4)
Qtz(33·1)Or(36)Ab(20·3)An(4·1)
Qtz(33·9)Or(36·1)Ab(19·5)An(4·7)
Qtz(32·5)Or(35·8)Ab(19·6)An(4)
Qtz(32·5)Or(35·4)Ab(22)An(3·8)
Qtz(34·9)Or(35·3)Ab(19·6)An(4·1)
Qtz(34·7)Or(34·1)Ab(20·1)An(3·9)
Qtz(32·5)Or(35·4)Ab(21·7)An(3·8)
Qtz(34·8)Or(34·7)Ab(19·2)An(3·8)
Qtz(38·9)Or(33·5)Ab(18·4)An(3·5)
Qtz(38·9)Or(34·4)Ab(16·9)An(3·5)
Qtz(38·2)Or(33·3)Ab(18·7)An(3·8)
Qtz(38·3)Or(33·3)Ab(19·4)An(2·9)
Qtz(36·2)Or(34·8)Ab(19·3)An(4·1)
Qtz(33·7)Or(32·6)Ab(22·4)An(4·4)
Qtz(34·6)Or(33·1)Ab(23·3)An(3·2)
Qtz(35·4)Or(34·1)Ab(20·5)An(3·9)
Qtz(35·9)Or(31·7)Ab(19·6)An(3·5)
Gl
X Usp(0·2)
Fo(6·5)Fa(91·6)
Fo(7)Fa(90·6)
Fo(9·3)Fa(88·1)
Fo(2·3)Fa(95·7)
Fo(9·5)Fa(88·5)
Fa
Fa
X Usp(0·3)
Mt*
X Usp(0·2)
X Usp(0·3)
X Usp(0·2)
Mt
X Usp(0·2)
X Usp(0·3)
X Usp(0·2)
Mt*
X Usp(0·2)
X Usp(0·2)
X Usp(0·2)
Mt
Mt
X Usp(0·5)
En(20)Wo(44·6)Fs(35·4)
BJR-57
An(3·5)Or(66·2)Ab(22·7)Cn(7·6)
An(4)Or(64·5)Ab(25)Cn(6·5)
An(4·3)Or(65·7)Ab(24)Cn(6)
An(2·4)Or(66·7)Ab(23·5)Cn(7·4)
An(4·7)Or(62·1)Ab(28·3)Cn(4·9)
An(6·4)Or(62·8)Ab(29·2)Cn(1·6)
An(2·4)Or(66·8)Ab(24·2)Cn(6·6)
An(5·1)Or(62·8)Ab(29·9)Cn(2·2)
Fsp*
An(2·2)Or(69·3)Ab(19·4)Cn(9·2)
An(3·1)Or(67·9)Ab(23·7)Cn(5·3)
An(3·6)Or(67·6)Ab(26·6)Cn(2·2)
Fsp*
An(2·3)Or(68·2)Ab(24·6)Cn(4·9)
An(3·5)Or(62·8)Ab(33·6)Cn(0·1)
An(3·5)Or(66·8)Ab(27·6)Cn(2·1)
An(2·7)Or(71·8)Ab(20·8)Cn(4·7)
An(6·4)Or(63·9)Ab(28·8)Cn(0·9)
Fsp
BJR-75
En(23·9)Wo(37·5)Fs(38·6)
BJR-54
En(19·1)Wo(15·7)Fs(65·2)
En(29·4)Wo(8·9)Fs(61·7)
An(30·8)Or(12·9)Ab(56·3)Cn(0)
Pl
Pl
NUMBER 9
X Usp(0·2)
Aug*
En(15·9)Wo(16·1)Fs(68)
En(19·3)Wo(35·6)Fs(45·1)
En(13·4)Wo(13)Fs(73·6)
En(9·5)Wo(9·1)Fs(81·4)
En(11·5)Wo(14)Fs(74·5)
Pig
VOLUME 53
BJR-51
En(26)Wo(38·2)Fs(35·8)
BJR-56
BJR-22
BJR-55
En(25·3)Wo(43·3)Fs(31·3)
BJR-73
En(32·4)Wo(42·3)Fs(25·3)
En(21·6)Wo(37·1)Fs(41·3)
BJR-72
BJR-58
En(17·5)Wo(32·5)Fs(50)
BJR-66
En(11·5)Wo(46·4)Fs(42·1)
En(26·3)Wo(47·8)Fs(25·9)
BJR-65
BJR-59
En(21·1)Wo(38)Fs(40·8)
BJR-64
En(19·7)Wo(37·8)Fs(42·5)
En(20·7)Wo(38·9)Fs(40·4)
BJR-62
BJR-60
En(14·2)Wo(31·1)Fs(54·7)
BJR-4
Aug*
En(18·3)Wo(24·7)Fs(57·1)
BJR-69
En(34·6)Wo(45·3)Fs(20·2)
En(26·2)Wo(43·1)Fs(30·7)
BJR-70
BJR-61
Aug*
BJR-71
BJR-23
Aug*
En(26·8)Wo(45·9)Fs(27·3)
BJR-39
Aug
Run
Table 4: Experimental runs and phase endmembers for Cougar Point Tuff rhyolite (BJR)
JOURNAL OF PETROLOGY
SEPTEMBER 2012
1849
En(33·3)Wo(45·1)Fs(21·6)
En(38·4)Wo(40·2)Fs(21·4)
En(46·2)Wo(39·4)Fs(14·4)
En(31·9)Wo(35·2)Fs(32·9)
IBR2-11
IBR2-13
IBR2-12
IBR2-53
En(34·9)Wo(17·2)Fs(47·8)
En(52·4)Wo(11·2)Fs(36·5)
En(36·8)Wo(21·8)Fs(41·4)
En(31)Wo(13·2)Fs(55·8)
Aug*
En(40·3)Wo(33·7)Fs(26)
IBR2-52r
IBR2-52
IBR2-53r
IBR2-70
Aug*
Aug*
IBR2-51
En(34·8)Wo(13)Fs(52·3)
En(34·5)Wo(13·5)Fs(52)
En(30·4)Wo(40·6)Fs(29)
IBR2-10
En(20·3)Wo(20·4)Fs(59·3)
En(33·7)Wo(12·6)Fs(53·7)
En(33·5)Wo(44·1)Fs(22·4)
En(39)Wo(11·6)Fs(49·4)
IBR2-50
En(32·7)Wo(37)Fs(30·3)
IBR2-35
En(42·1)Wo(41·1)Fs(16·8)
IBR2-16
IBR2-34
En(31·5)Wo(38·7)Fs(29·9)
IBR2-15
En(29·1)Wo(27·6)Fs(43·4)
En(32·6)Wo(42)Fs(25·5)
IBR2-18
IBR2-14
En(38·6)Wo(46·3)Fs(15·2)
IBR2-26
En(42·7)Wo(44·3)Fs(13)
En(40·5)Wo(42·4)Fs(17·1)
IBR2-28
IBR2-17
En(25·4)Wo(37·4)Fs(37·3)
En(38·2)Wo(42)Fs(19·7)
IBR2-27
An(41)Or(10·5)Ab(48·3)Cn(0·2)
An(34·6)Or(14)Ab(51·4)Cn(0)
An(35·1)Or(13·5)Ab(51·4)Cn(0)
An(30·4)Or(17·4)Ab(51·6)Cn(0·7)
An(31·9)Or(17·6)Ab(50)Cn(0·5)
An(35·2)Or(9·7)Ab(54·8)Cn(0·2)
An(38·2)Or(11)Ab(50·6)Cn(0·2)
An(27·4)Or(20·3)Ab(51·7)Cn(0·6)
An(30·4)Or(13·6)Ab(56)Cn(0)
Pl*
Pl*
An(35·5)Or(17·6)Ab(46·8)Cn(0·1)
Pl*
An(24·5)Or(23·1)Ab(51·5)Cn(0·9)
An(40·7)Or(10)Ab(49·1)Cn(0·3)
An(37·5)Or(9·2)Ab(53·1)Cn(0·2)
An(35·4)Or(11·7)Ab(52·6)Cn(0·3)
En(32·1)Wo(35·3)Fs(32·6)
IBR2-25
IBR2-24
En(34·8)Wo(33·7)Fs(31·5)
IBR2-21
An(6·4)Or(57·2)Ab(34·3)Cn(2·1)
An(6·2)Or(56·7)Ab(33·3)Cn(3·8)
An(6·8)Or(58)Ab(34·3)Cn(0·8)
An(4·5)Or(63·5)Ab(31·5)Cn(0·5)
An(4·7)Or(61·8)Ab(30·7)Cn(2·8)
An(4·7)Or(61·6)Ab(33)Cn(0·6)
An(3·5)Or(67·2)Ab(28·2)Cn(1·1)
An(8·5)Or(54·1)Ab(36·5)Cn(0·9)
Fsp
An(7·9)Or(60·9)Ab(30·9)Cn(0·3)
Pl
En(37·2)Wo(33·7)Fs(29·2)
Pig
IBR2-20
Aug
IBR2-23
Run
Table 5: Experimental runs and phase endmembers for Indian Batt rhyolite (IBR2)
Mt*
X Usp(0·3)
Mt*
X Usp(0·4)
X Usp(0·3)
X Usp(0·6)
Mt*
X Usp(0·3)
Mt*
(continued)
Qtz(25·7)Or(28·6)Ab(22·7)An(6·9)
Qtz(26·9)Or(30·6)Ab(24·2)An(7·2)
Qtz(32·2)Or(33·3)Ab(22)An(5·5)
Qtz(29·5)Or(33·4)Ab(24·3)An(4·8)
Qtz(29·3)Or(33·8)Ab(23·8)An(4·6)
Qtz(36·8)Or(33·0)Ab(15·0)An(5·5)
Qtz(30·3)Or(35·5)Ab(22·4)An(5·3)
Qtz(30·5)Or(35·4)Ab(21·9)An(5·9)
Qtz(32·2)Or(34·8)Ab(22·7)An(4·6)
Qtz(34·8)Or(35·7)Ab(20·8)An(3·8)
Qtz(31·6)Or(36·2)Ab(22·1)An(3·8)
Qtz(28·5)Or(32·2)Ab(25·2)An(3·9)
Qtz(26·6)Or(34·8)Ab(24·8)An(3·8)
Qtz(29·7)Or(31·3)Ab(23·5)An(8·4)
Qtz(27·9)Or(36·9)Ab(23·8)An(2·4)
Qtz(32·4)Or(33·7)Ab(23)An(4·7)
Qtz(36·2)Or(35·6)Ab(19·7)An(3·1)
Qtz(34·9)Or(36·2)Ab(20·6)An(3)
Qtz(31·2)Or(30·6)Ab(21·6)An(8)
Qtz(31·8)Or(32·6)Ab(22·2)An(5·7)
Qtz(33·8)Or(33·7)Ab(20·9)An(4·4)
Qtz(32·8)Or(34·8)Ab(22·4)An(4·6)
Qtz(35·5)Or(35·7)Ab(20·2)An(3·3)
Qtz(37·5)Or(33·5)Ab(19·5)An(3·1)
Qtz(36·1)Or(35·8)Ab(19·5)An(3·1)
Qtz(35·9)Or(33·2)Ab(21)An(3·5)
Gl
Mt*
Fa
X Usp(0·2)
X Usp(0·3)
X Usp(0·5)
X Usp(0·2)
X Usp(0·3)
X Usp(0·3)
X Usp(0·3)
Mt*
X Usp(0·4)
X Usp(0·4)
X Usp(0·3)
Mt
ALMEEV et al.
SILICIC MAGMAS OF YELLOWSTONE HOTSPOT
En(26·1)Wo(17·3)Fs(56·6)
En(42·9)Wo(9·6)Fs(47·5)
An(24·1)Or(27·2)Ab(48)Cn(0·7)
An(37·9)Or(14·7)Ab(47·4)Cn(0)
An(10·1)Or(48·5)Ab(40·7)Cn(0·7)
X Usp(0·5)
Qtz(26·3)Or(29·3)Ab(26·4)An(7·3)
Qtz(25·8)Or(30·2)Ab(26·8)An(4·8)
Qtz(28·2)Or(32·4)Ab(24·6)An(6)
Qtz(28·7)Or(32·7)Ab(24)An(6·4)
Qtz(27·3)Or(33·5)Ab(25·5)An(5·2)
1850
En(34)Wo(29·8)Fs(36·2)
IBR2-63
*Detected phases without microprobe measurements and phase calculations.
Qtz(25·2)Or(28)Ab(23·6)An(6·8)
Qtz(25·9)Or(29·8)Ab(25·5)An(7·1)
NUMBER 9
IBR2-61
En(41·2)Wo(34·2)Fs(24·7)
IBR2-60
Qtz(25·9)Or(30·1)Ab(26)An(6·2)
En(37·3)Wo(26·6)Fs(36·1)
IBR2-59
En(40)Wo(19)Fs(41)
Qtz(27)Or(29·7)Ab(24·6)An(6·8)
Qtz(28·5)Or(34)Ab(23·3)An(5·1)
IBR2-58
En(32·7)Wo(15·6)Fs(51·6)
Aug*
IBR2-65
Qtz(25·6)Or(30·6)Ab(25·4)An(6·6)
Qtz(26·2)Or(30·3)Ab(24·5)An(6·9)
Qtz(26·9)Or(34·3)Ab(24·5)An(4·3)
En(37·7)Wo(24·9)Fs(37·4)
IBR2-64
VOLUME 53
En(40·4)Wo(16·1)Fs(43·5)
Qtz(26·2)Or(27·5)Ab(20)An(7·9)
An(8)Or(54·4)Ab(36·8)Cn(0·8)
Qtz(27·7)Or(30·2)Ab(23·5)An(7·6)
Aug*
En(28·3)Wo(15·4)Fs(56·3)
Qtz(26·6)Or(30·4)Ab(24·8)An(7)
Qtz(26·4)Or(29·8)Ab(25·2)An(7)
IBR2-62
En(39·3)Wo(35·8)Fs(24·9)
IBR2-69
En(36·9)Wo(14·8)Fs(48·3)
IBR2-71
En(33·1)Wo(28·6)Fs(38·4)
IBR2-68
Qtz(27·7)Or(34·1)Ab(24·2)An(4·8)
En(44·6)Wo(41·3)Fs(14)
IBR2-48
An(36·7)Or(18)Ab(44·8)Cn(0·5)
X Usp(0·6)
Qtz(28·5)Or(34)Ab(23·9)An(5·8)
Gl
En(29·8)Wo(27·5)Fs(42·8)
Aug*
IBR2-38
En(48·7)Wo(8·8)Fs(42·5)
An(44·8)Or(11·5)Ab(43·5)Cn(0·2)
Fa
IBR2-67
En(42·9)Wo(40)Fs(17·1)
IBR2-31
En(46·5)Wo(8·8)Fs(44·6)
An(39·4)Or(12·6)Ab(47·6)Cn(0·4)
An(36·8)Or(16·5)Ab(46·4)Cn(0·3)
Mt
Qtz(25)Or(28·7)Ab(24·2)An(6·7)
En(38·1)Wo(36·5)Fs(25·4)
IBR2-32
En(40·1)Wo(9·8)Fs(50·1)
En(41·3)Wo(9·9)Fs(48·8)
Fsp
Aug*
Aug*
IBR2-30
Pl
IBR2-66
En(37·1)Wo(30·6)Fs(32·3)
IBR2-29
Pig
IBR2-49
Aug
Run
Table 5: Continued
JOURNAL OF PETROLOGY
SEPTEMBER 2012
ALMEEV et al.
SILICIC MAGMAS OF YELLOWSTONE HOTSPOT
Fig. 3. Back-scattered electron images of run products obtained at 500 MPa and 9758C on IBR2 with different water contents in the glass
(7·3 wt %, 3·0 wt %, 2·3 wt %, 2·0 wt % and 1·0 wt % H2O).
Fig. 4. Phase relations determined for the composition BJR at 200 MPa as a function of temperature and the water content in the melt
(in wt %). The experimental data are given in Table 2. The solidus curve is estimated from the synthetic Qtz^Ab^Or system. Black arrows
show different redox conditions relative to QFM buffer. Mineral abbreviations are as in Fig. 1. In the field labeled ‘melt (þ Ox)’, either pure
melt or melt coexisting with an oxide phase (mostly magnetite, see Tables 2 and 3) was observed.
experiments (1025 and 9758C, H2Omelt51wt %), magnetite is not observed but ilmenite has been detected. The
similar phase assemblage observed at oxidizing and reducing conditions, independently of fO2, is attributed to the
low Fe contents in the bulk samples. Even if the
proportions of magnetite vary as a function of fO2, or
even if ilmenite is present instead of magnetite, the abundance of oxide phases is so low that an effect on the composition of the silicate melt and on the crystallization
sequence of the other phases is expected to be minor.
1851
JOURNAL OF PETROLOGY
VOLUME 53
NUMBER 9
SEPTEMBER 2012
Fig. 5. Phase relations determined for the composition IBR2 at 200 MPa as a function of temperature and the water content in the melt
(in wt %). The experimental data are given in Table 3. (For other details see the caption to Fig. 4.)
IBR2 (500 MPa)
A limited number of experiments were conducted at
500 MPa at temperatures between 975 and 10258C and
H2Omelt between 0·7 and 7·3 wt %. These data are sufficient to estimate at least qualitatively the effect of pressure
on the stability curves. Compared with the 200 MPa
phase stability fields, the Cpx-in curve and the Pig-in
curve are shifted to higher temperatures (Fig. 6). The
Qtz-in and San-in curves are also shifted to higher temperatures (by 50^708C) but the position of the Plag-in
curve does not change significantly when compared with
the 200 MPa curve. It is emphasized that the Plag-in
curve is constrained by only three experiments (IBR2-66,
IBR2-62, IBR2-58; black squares in Fig. 6) in which single
plagioclases could not be resolved and analyzed as stoichiometric homogeneous phases. In one experiment
(IBR2-58), analysis of the feldspar phase can only be interpreted as a mixture of sanidine and plagioclase (An24^
Or27^Ab48^Cn1; Table 5). Such an intergrowth may result
from strong undercooling (simultaneous crystallization of
both feldspar phases). In the other two experiments,
plagioclase could not be clearly identified because it
occurs as small intergrowths within the highly crystalline
parts of the run products (see IBR2-66 in Fig. 3).
However, plagioclase must have crystallized in these
experiments, because the Ab content in the glass is
significantly lower than that of the starting material (see
discussion below).
Phase compositions
Plagioclase
Plagioclase was produced almost exclusively in the less
evolved (more mafic) composition IBR2. The composition
of plagioclase synthesized in the IBR2 experiments is difficult to assess because of the difficulty in analyzing single
plagioclase grains in highly crystalline run products. In
several experiments with high crystallinity, the analyses of
plagioclase yield very high Or contents (Or13^23, Table 5),
probably as a result either of the presence of small melt inclusions or of intergrowth of sanidine and plagioclase. If
only experiments in which plagioclase can be accurately
analyzed (Or contents of plagioclase 5Or11) are taken into
account, the An content increases with increasing
H2Omelt, as expected from previous studies (Klimm et al.,
2003; Lange et al., 2009). For example, at 200 MPa the An
content increases from 35 to 41 with increasing H2Omelt
from 1·3 to 2·8 wt % at 9008C; the An content increases
from 39 to 45 with increasing H2Omelt from 0·5 to 0·7 wt
% at 10008C (Table 5, composition IBR2). The composition of the experimentally produced plagioclases with Or
contents below 5Or11 overlaps with the natural compositional range in the rhyolite lava of An33^40 in all
1852
ALMEEV et al.
SILICIC MAGMAS OF YELLOWSTONE HOTSPOT
Fig. 6. Phase relations determined for the composition IBR2 at 500 MPa as a function of temperature and the water content in the melt (in wt
%). The experimental data are given inTable 3. The grey lines are stability fields at 200 MPa. The curvature of the pigeonite-in boundary is estimated from the lower pressure experiments. (For other details see the caption in Fig. 4.)
experimental products for IBR2 (except for 10008C, 0·5 wt
% H2Omelt). However, because of the difficulty in analyzing single plagioclase crystals in the experiments at very
low water activity, the An content is difficult to use to accurately constrain the H2Omelt prevailing in silicate melts
in equilibrium with natural plagioclase.
below). The Or content in experimental sanidine produced
with the IBR2 composition varies in the range Or48^68,
and Cn contents are generally lower than in BJR, with
maximum values of 3·8 wt %. Sanidine has not been
found in the natural sample of the Indian Batt rhyolite
and cannot be compared with the synthetic sanidine.
Sanidine
Fayalite
The Or contents in experimental sanidine produced with
the BJR composition vary in the range Or62^68 for sanidine
with low celsian (Cn) contents (52·2 wt %). This Or content is slightly higher than that analyzed in natural sanidines. In the BJR rock sample, the sanidine composition
is in the range Or49^60 (this study) with Cn of 0^4 wt %.
Interestingly, the Cn content is relatively high in some experiments (up 9 wt % in BJR) owing to the high Ba content of the bulk-rock composition (Table 1). At a given
temperature, the Cn content in sanidine increases with
increasing H2Omelt (at 8758C, Cn increases from 2·2 to
9·2 with increasing H2Omelt from 1·1 to 2·2 wt %; Table 2)
which may be due to the high partition coefficient of Ba between sanidine and melt (Leeman & Phelps, 1981; Nash &
Crecraft, 1985; Icenhower & London, 1996) and to the decrease in the abundance of sanidine in the experimental
products with increasing H2Omelt (see also discussion
Fayalite is observed in six experiments with high crystal
contents (close to the solidus). Fayalite was difficult to analyze by EPMA and the results should be interpreted with
caution. However, except for one experiment at 10008C
(BJR-75) the experimental composition is close to that of
the natural phase (Fo12Fa88).
Pyroxenes
BJR. The detailed analysis of the composition of pyroxenes
(Fig. 7) confirms that some experimentally produced augites with Wo contents higher than 40% do not overlap the
natural augite range. These clinopyroxenes were synthesized at relatively high H2Omelt (41wt % H2O at 900 and
9258C; 42 wt % H2O at 8758C; 43 wt % H2O at 8258C;
filled triangles in Fig. 7a). Augites synthesized in melts
with lower H2Omelt have compositions that overlap the
natural augite range (open triangles in Fig. 7a). These are
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augites synthesized at 9758C and 9508C (0·3^0·6 wt %
H2O), 9258C (0·8 wt % H2O), 9008C and 8758C (0·9 to
1·8 wt % H2O). Most of the experiments performed at
these conditions also contain pigeonite. In general, the
composition of augite synthesized experimentally at low
water contents falls within the range defined by the natural
CPT augite (En9^32Wo32^41Fs31^52) although some experimental augites trend toward slightly higher Fs contents.
Pigeonites in BJR experiments fall along the natural
pigeonite trend, but do not reproduce the more En-rich
compositions observed in the CPT, and extend to more
Fe-rich compositions (i.e. XFs471). It should be emphasized that in the natural assemblages, zoning in pyroxene
is rare; the vast majority of pyroxenes are unzoned
(Cathey & Nash, 2004).
IBR2. For clarity and for a better comparison with the
BJR data, only augites obtained at 200 MPa (intrinsic fO2
conditions) are plotted in Fig. 7b. The experimentally produced augites that depart significantly in composition
from those in the natural rocks are Mg- and Ca-rich and
Fe-poor (black squares in Fig. 7b) compared with the
observed augite compositions of the natural rock. They
also coexist with more Mg-rich pigeonite than observed
in both lavas and in members of the CPT. These Mg-rich
pyroxenes were synthesized at relatively high H2Omelt
(40·8 wt % H2O at 10008C; 41·5 wt % H2O at 9508C;
41·7 wt % H2O at 9258C; 42·4 wt % H2O at 9008C) and
high fO2. The experimental pyroxenes (augite and coexisting pigeonite) that have compositions similar to those in
the natural IBR2 sample were synthesized in experiments
with lower H2Omelt (open squares in Fig. 7b) and consequently lower fO2. These are augite and pigeonite synthesized at 10008C (0·5^0·7 wt % H2O), 9758C (0·5 wt %
H2O), 9508C (0·5^1·2 wt % H2O), 9258C (0·8^1·0 wt %
H2O) and augite only at 9008C (1·1wt % H2O) and
8758C (1·4^1·8 wt % H2O).
Collectively, the results reproduce and extend beyond
the range of pyroxene compositions observed in the
Cougar Point Tuff and Bruneau^Jarbidge eruptive center
lavas, with the more mafic IBR2 reaching more En-rich
and the more evolved BJR reaching more Fs-rich compositions. The spectrum of pyroxene compositions synthesized
from IBR2 overlaps with the Mg-rich end of the range
produced from BJR, but reaches higher En content, consistent with its less evolved composition. Figure 8 summarizes the variation of En content in pyroxenes as a function
of temperature and wt % H2O for all experiments (see
data in Table 4 and Supplementary File SM-1). There is a
strong positive correlation between En mol % in pigeonite
and temperature (especially evident in the BJR sample)
that is similar to that observed in the CPT (Cathey &
Nash, 2004, 2009), and a less well-defined but overall positive correlation between En content in augite and temperature. In addition, all low-temperature experiments are
NUMBER 9
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conducted at high aH2O, and therefore high fO2 (see control of fO2 in experiments), which explains the general increase in mg# with increasing fO2 (see figures in
Supplementary File SM-2). Taken collectively, the data suggest that conditions that best reproduce the observed compositions of pigeonite^augite pairs in the starting natural
samples are 0·4^1·0 wt % H2O and 975^9008C for the
BJR sample (200 MPa), and 0·5^1·3 wt % H2O and 1000^
9258C for the IBR2 sample (200 and 500 MPa). The corresponding fO2 of most experiments in this range of temperatures and melt water concentrations is between QFM
and QFM þ1·5, which is realistic for natural conditions
(Christiansen & McCurry, 2008).
Glasses
The major element composition of the glasses in the experimental products is given in Supplementary File SM-1 and is
plotted in Figs 9 and 10 as oxides vs SiO2. There is little variation in the quartzo-feldspathic components (Na, K, Al
and Si) of the melt as a function of temperature or melt
water content, especially for the BJR composition, which is
to be expected considering that the bulk composition is
very close to the minimum or eutectic composition of the
granitic system. The FeO contents in the glass vary slightly
and, as expected from previous studies on granitic systems
(e.g. Scaillet et al., 1995, Scaillet & Evans, 1999; Klimm
et al., 2003), FeO concentrations decrease with decreasing
temperature and decreasing water activity (at a given temperature). For both BJR and IBR2, CaO, MgO, FeO, and
TiO2 decrease with increasing SiO2 content. The more
SiO2-rich compositions of the experimental glasses produced for BJR fall within the range of natural glasses from
the CPT (Fig. 9). The MgO, FeO, and TiO2 concentrations
of the experimental IBR2 glasses overlap with the compositional range observed for the whole-rock compositions of
the lavas and the more evolved CPT from the Bruneau^
Jarbidge eruptive center, except for some experiments with
lower FeO concentrations (Fig. 10). Some experimental
glasses obtained at the most oxidizing conditions have
slightly lower FeO concentrations than the natural trend,
which may be explained by the presence of magnetite.
The variations in Al2O3, Na2O, K2O, and to a lesser
extent CaO are related to the crystallization of tectosilicates, which are the most abundant phases in the experimental products. In both starting compositions, CaO
concentrations decrease with increasing SiO2. The evolution of Al2O3 as a function of SiO2 is not as systematic
and depends on the crystallization of plagioclase. The increase of Al2O3 with SiO2 in BJR and for some glasses in
IBR2 may be related to the absence of plagioclase in the
experimental products. The Na/K ratio in the glasses from
BJR is roughly constant, but decreases with increasing
SiO2 in the glasses from IBR2. This difference may also
be related to the absence of plagioclase in the BJR experimental products when compared with the IBR2 system.
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Fig. 7. Compositions of natural and synthetic pyroxenes (Cpx and Pig) in experimental products from BJR and IBR2. The data points are average values obtained for each experiment and have been normalized and re-plotted in the En^Fs^Wo triangular diagram. The dashed tie lines
denote mineral pairs obtained for the same experiment (equal temperature and H2Omelt). Black filled squares and triangles: compositions that
depart significantly from the natural compositions (mainly in experiments with high H2Omelt). Open squares and triangles: compositions that
are close to the natural compositions; mainly for experiments with low H2Omelt. Diamonds: compositions determined in the natural samples.
Circles: pyroxene compositions reported for other units or lavas by Cathey & Nash (2004, 2009).
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Fig. 8. Variation of En content (mol %) in pyroxenes vs temperature and H2Omelt in the experiments. Triangles represent results from experiments using the BJR (Cougar Point Tuff XVj) bulk composition at 200 MPa, and squares represent compositions from the IBR2 (Indian Batt
Rhyolite) bulk composition at 200 and 500 MPa. The field bounded by the white line indicates the range of En vs T determined for Cougar
Point Tuff pigeonite based on QUILF thermometry (Cathey & Nash, 2004; note that the projected compositions here include only the Wo^
En^Fs components). The ranges of observed compositions in all members of the Cougar Point Tuff and lavas are indicated by the grey bands
and vertical arrows at the right; the dashed horizontal lines indicate the natural compositions in the starting materials (CPT XVj is bimodal
with respect to pigeonite). The T^H2Omelt conditions that best reproduce the observed natural mineral compositions in the BJR and IBR samples are indicated by the shaded grey ovals (bottom of figures). (See text for discussion.)
The normative Qtz, Ab, Or and An contents of the
glasses are given in Tables 4 and 5 and plotted in ternary
Qtz^Ab^Or diagrams in Figs 11 and 12. The projected
compositions were corrected to account for the effect of
the An component, which is not present in the haplogranite projection using the equations given by Blundy &
Cashman (2001). The expressions proposed by those workers match the coordinates of cotectic compositions from
An-free systems and An-bearing systems (20% An)
when projected into the Qtz^Ab^Or diagram. For the
BJR composition, most of the experimental glasses obtained at 200 MPa (open and black triangles in Fig. 11)
have Qtz^Ab^Or proportions that are very similar to
those of the starting material. Four glasses synthesized at
825 and 8758C have slightly higher Qtz contents. The
glasses obtained from the IBR2 composition at 200 MPa
show a clear compositional variation (grey and black
squares, Fig. 12). The Ab content in the glasses shows continuous variation from nearly identical to that of the starting composition to 15% less. The Qtz/Or ratio remains
approximately constant in all 200 MPa IBR2 glasses. The
compositions of the experimental glasses obtained at
500 MPa with the starting material IBR2 do not differ
significantly from that of the starting material (grey
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SILICIC MAGMAS OF YELLOWSTONE HOTSPOT
Fig. 9. Binary diagrams showing the variation in composition of the experimental glasses (triangles) obtained for the BJR composition. The
crosses (þ) represent the composition of the natural groundmass glasses in BJR. Cougar Point Tuff glass analyses (grey dots) are from vitrophyres and air-fall tuffs (Cathey & Nash, 2004; Nash B.P., unpublished data). Circled star is the starting composition BJR. The experimental
glasses were normalized to 100%.
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Fig. 10. Binary diagrams showing the variation in composition of the experimental glasses (squares and diamonds) obtained for the IBR2 composition. The compositional fields of whole-rock compositions for the Cougar Point Tuffs and other rhyolite lavas (post Cougar Point Tuff
lavas) from the Bruneau^Jarbidge eruptive center are from Bonnichsen et al. (2008). The crosses () represent the compositions of the natural
groundmass glasses in the investigated sample. The circled stars are the starting compositions of BJR and less evolved IBR2 used in the
experiments.
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ALMEEV et al.
SILICIC MAGMAS OF YELLOWSTONE HOTSPOT
Fig. 11. Compositions of all the experimental glasses (triangles) obtained with the BJR composition (star) plotted in the Qtz^Ab^Or system.
The Qtz^Ab^Or proportions have been calculated on the basis of the CIPW norm and are corrected for the An content following the procedure
of Blundy & Cashman (2001). Filled triangles are glasses from experiments in which melt coexists with quartz, sanidine and sometimes plagioclase. Open triangles are glasses from the other experiments. The black diamonds correspond to water-saturated minimum and eutectic compositions at 200 MPa (Tuttle & Bowen, 1958; Qtz36Ab39Or25), 500 MPa (Holtz et al., 1992; Qtz31Ab47Or22) and 1GPa (Luth et al., 1964;
Qtz23Ab56Or21). Black circles correspond to water-undersaturated minima at 200 and 500 MPa for water activities of 0·25 and 0·4, respectively
(Holtz et al.,1992; Qtz36Ab33·5Or30·5 at 200 MPa; Qtz31Ab40·5Or28·5 at 500 MPa). Grey circles are analyses of natural CPTglass from vitrophyres
and air-fall tuffs (Cathey & Nash, 2004; Nash B.P., unpublished data). The crosses (þ) represent the composition of the natural groundmass
glasses in the investigated sample.
diamonds, Fig. 12), except for two glasses that may coexist
with plagioclase and have higher Qtz and Or contents
(black diamonds, Fig. 12). Experimental glasses from
products containing at least two tectosilicate phases have
been distinguished from the other glasses in Figs 11 and
12. Although the compositional differences are small,
glasses with cotectic compositions in IBR (experiments
with quartz, sanidine, and sometimes plagioclase) tend
to have the highest Qtz and Or contents (at the expense
of Ab).
DISCUSSION
Attainment of equilibrium
Experiments with different run durations at the same P^T
conditions and aH2O were not performed in this study.
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Fig. 12. Compositions of all experimental glasses at 200 and 500 MPa (squares and diamonds, respectively) in the IBR2 composition (star)
plotted in the Qtz^Ab^Or system. Filled symbols are glasses from experiments in which melt coexists with quartz, sanidine and sometimes
plagioclase. Grey symbols are glasses from the other experiments. Crosses () represent the compositions of the natural groundmass glasses
in the investigated sample.
However, the run durations (Tables 2 and 3) are estimated
to be sufficient to attain near-equilibrium conditions in
such highly polymerized melt compositions, at least in the
high-temperature experiments with H2Omelt of 1^2 wt %.
This is confirmed by the investigation of the kinetics of
crystallization in hydrous rhyolitic melts (3^7 wt % H2O)
in experiments with similar starting materials (dry glass
powder þ added fluid) conducted at relatively low temperature by Pichavant (1987) and Holtz et al. (1992). Those
workers demonstrated that the use of crushed glasses is
crucial to reach equilibrium conditions because
heterogeneities at grain boundaries favour the nucleation
of crystals (see Fig. 3; H2Omelt ¼1·0 wt %). High nucleation rates result in numerous small phases that are difficult to analyze (with intergrowth of phases such as
sanidine and plagioclase), but reduce considerably the
transport distances for those elements with low diffusivities
such as Si and Al. Experiments in rhyolitic systems conducted with seeds of either quartz or feldspar have confirmed that equilibrium phase assemblages can be
determined using the approach described above
(Pichavant, 1987; Holtz et al., 1992). Crystallization
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ALMEEV et al.
SILICIC MAGMAS OF YELLOWSTONE HOTSPOT
experiments from powdered starting material conducted
with nearly dry rhyolitic melt compositions (51wt %
H2O) at high temperature (1020^11008C) by Becker et al.
(1998) and in completely dry systems at 1atm by Brugger
et al. (2003) at 10508C also indicate that run durations of
more than 160 h are sufficient to reach near-equilibrium
conditions, at least for experiments at 9008C and above.
Brugger et al. (2003) demonstrated that near-equilibrium
conditions could be reached at 10508C in a dry rhyolitic
composition containing 78·1wt % SiO2 and 12·8 wt %
Al2O3 after 82 h. At this temperature (10508C), the viscosity of such melts, which controls directly the diffusivity of
the network-forming cations, is 107·2 Pa s [following the
model of Giordano et al. (2008)]. Thus, if the viscosity of
the silicate melts investigated in this study is lower than
107 Pa s, the chosen experimental duration (160 h or more)
should be sufficient to approach near-equilibrium conditions. The viscosity of the BJR silicate melt at 9008C with
0·75 wt % H2O is 106·8 Pa s, and this value corresponds to
the highest viscosity in our experiments at 9008C, considering that the experimental melts at this temperature
always contain at least 0·75 wt % H2O. Similarly, the highest viscosity of the BJR melt obtained in this study at
9508C is 106·8 Pa s (viscosity of melts with 0·4 wt % H2O).
Thus, even if the distribution of the minerals is heterogeneous at the scale of 100 mm as a result of the structure of
the starting material (powder of dry glass grains) in the
experiments with high crystallinity, information derived
from kinetic studies and melt viscosity indicates that
near-equilibrium conditions should be reached in most of
the experiments.
To check for possible disequilibrium features, the glasses
in experiments with a heterogeneous distribution of
phases have been analyzed in detail for selected experiments. Assuming that equilibrium is not reached between
the core and the rim of the initial glass grains (where
most of the crystallization takes place), different compositions are expected to be observed between the core and
the rim. In all cases, the glass composition was homogeneous (within analytical uncertainty), even for
network-forming cations. For example, in experiment
IBR2-58 (500 MPa, 10258C, 0·7 wt % H2O), the SiO2 content in the starting dry glass is 70·7 wt %, and the SiO2
contents in the core and rim of melt pools are 72·4 and
72·5 wt %, respectively. Another example yielded values
of 71·8 and 72·1wt %, respectively. The good agreement
between the compositions of melts coexisting with quartz
and sanidine and the position of the cotectic line in the
Qtz^Ab^Or diagram (see above and Figs 11 and 12) also
indicates that near-equilibrium conditions were reached.
An additional test to check for the attainment of
near-equilibrium conditions in the experiments is the analysis of trace element concentrations in the experimental
phases. The relatively high Ba concentration in the starting
material yields good data on the distribution of Ba between
alkali feldspar and melt. In experimental products
that were analyzed with accurate EPMA conditions for
Ba (40 nA), the distribution of Ba between alkali feldspars
and glasses, D(Ba)Fsp/Gl is in the range of 9^18 (mean
value of 24 experiments is 14 2·5). Such values are consistent with the broad range reported in the literature (Nash
& Crecraft, 1985; Ren, 2004), and are very close to the
D(Ba)Fsp/Gl of 16 0·5 expected from the study of
Icenhower & London (1996).
Magma storage conditions for the samples
BJR and IBR2
Although the experimental pressure of 200 MPa may not
correspond exactly to the depth of the magma reservoir
prior to eruption of the Indian Batt Rhyolite, the mineral
assemblage and mineral compositions of the mafic phases
in IBR2 can be used to constrain the temperature and
H2Omelt. The mineral assemblage and compositions of the
phases observed in the natural sample (Aug, Pig, Plag; absence of Qtz and San) can be reproduced at 200 MPa in a
narrow T^H2Omelt range at temperatures higher than
9208C. The mineral assemblage is reproduced if H2Omelt
is in the range 0·7^1·3 wt % at 9508C and 0·3^1·0 wt % at
10008C (Fig. 5). This range is also within the T^H2Omelt
range at which the compositions of the synthetic augite
and pigeonite reproduce the compositions of the natural
phases (Fig. 7b). QUILF thermometry (Andersen et al.,
1993) for the Indian Batt Rhyolite using pigeonite yields a
temperature of 975 118C (Cathey & Nash, 2009) that
would indicate an H2Omelt of 1·0 wt %.
The experimental dataset at 500 MPa is more limited
but quartz, sanidine and probably plagioclase start to crystallize within a very similar T^H2Omelt range. Although
the stability field of the tectosilicate phases in this low
T^H2Omelt range needs to be worked out in more detail
at 500 MPa, the phase diagram in Fig. 6 allows us to determine the boundary conditions at which the three tectosilicate phases may be present. Plagioclase crystallization
may start at the same temperature as sanidine and quartz
at H2Omelt51wt % at 10258C and H2Omelt51·5 wt % at
9758C. At such conditions pigeonite and clinopyroxene are
also stable.
The comparison of the phase equilibria at 500 and 200
MPa indicates that pressure has a significant effect on the
stability of tectosilicates, and this observation can be used
to constrain the depth of magma storage prior to eruption.
Assuming that the natural plagioclase phenocrysts are in
equilibrium with the silicate melts (there is no evidence
for dissolution or reaction rims around most natural
plagioclase phenocrysts), and taking into account that
quartz and sanidine are not present in the natural assemblage, magma storage pressures for the IBR2 sample
should be less than 500 MPa. Clearly, the effect of pressure
on plagioclase, quartz and feldspar stability needs to be
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JOURNAL OF PETROLOGY
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better calibrated for appropriate rhyolitic compositions at
various pressures; however, the dataset obtained so far indicates that a pressure higher than 500 MPa is not realistic
for the magma storage conditions of sample IBR2.
The effect of pressure on phase relationships was not
investigated for BJR, and the results at 200 MPa are
used to discuss the possible T^H2Omelt range of the BJR
magma prior to eruption. Fayalitic olivine, augite,
pigeonite, quartz and sanidine coexist in some experimental products. These phases are also observed in the
natural sample, and taking into account that the glass
proportion is at least 80%, the pre-eruptive conditions
can be constrained to a field in which H2Omelt is
51·5 wt %. This maximum value is governed by the absence of pigeonite at higher H2Omelt. Augite and pigeonite compositions are strongly dependent on temperature
and H2Omelt, and compositional pairs similar to those
in the natural system (Fig. 7) could be reproduced only
at low water contents (52·2 wt % at 875 wt %; 51·0 wt
% at 9258C; 50·8 at 9758C; Fig. 7a). In addition, the
Fs content of pyroxenes from the augite^pigeonite pair
obtained in the experiment at the lowest temperature
(8758C, H2Omelt ¼1·1wt %) is higher than that of the
natural pairs, and this T^H2Omelt combination is probably not realistic. Temperature estimations using pyroxene geothermometers yield values of 9508C (Cathey &
Nash, 2004) and the natural assemblage (neglecting
plagioclase) can be reproduced at such temperatures
only if H2Omelt is very low (50·9 wt % H2O). It is
worth noting that plagioclase was almost never observed
in the experimental products (except two runs at 825
and 8508C) but this mineral is present as phenocrysts
in the natural sample. The absence of plagioclase in the
experimental products is related to the relatively low
Al2O3 concentration (11·8 wt %) in the BJR starting
composition when compared with other samples from
the Bruneau^Jarbidge eruptive center. A possible explanation for this compositional feature is given below.
The discussion above assumes that the mineral assemblages and glasses in sample BJR are in equilibrium.
However, there is evidence that this may not completely
be the case and that the phenocrysts and the silicate
melt in the BJR sample may come from a compositionally heterogeneous magma reservoir. Cathey & Nash
(2004) considered this to be a possibility because single
hand samples of members of the Cougar Point Tuff commonly contain multiple modes of augite and pigeonite,
indicating derivation from slightly different regions of
the magma reservoir and mixing upon eruption. An alternative explanation for the discrepancy between the experimental and the natural phase assemblages may be
related to the pressure at which the experiments were
conducted. An indication that some minerals may be inherited from deeper magmatic reservoirs is given by the
NUMBER 9
SEPTEMBER 2012
composition of the natural sanidines, which could not
be perfectly reproduced. The natural phases have systematically slightly lower Or contents than the minerals
synthesized at 200 MPa. This observation, combined
with data on the composition of feldspars coexisting
with rhyolitic or haplogranitic melts (e.g. Tuttle &
Bowen, 1958; Holtz et al., 1992), suggests that such
Ab-rich sanidine may have crystallized at a pressure
higher than 200 MPa. At higher pressure in the system
Qtz^Ab^Or, the feldspar solvus is close to the solidus,
leading to a shift in the feldspar composition towards
the end-member compositions (either Ab-rich or
Or-rich), and intermediate compositions are not expected
to crystallize at equilibrium conditions.
Differentiation processes in the
Bruneau^Jarbidge eruptive center
A comparison of the composition of the glasses obtained
in the experiments with natural rock samples and glasses
from the Cougar Point Tuff rhyolites and the Bruneau^
Jarbidge eruptive center in general is useful to interpret
the compositional trends in the natural systems (Figs 9
and 10). The starting materials BJR and IBR2 are
shown in Figs 9 and 10; these compositions represent
intermediate-silica and low-silica rhyolites, respectively,
within the range of rhyolite compositions observed from
the Bruneau^Jarbidge eruptive center. Figure 10 shows
that the general trends observed in the natural BJR
rhyolites (including the CPT tuffs) can be reproduced
experimentally using the IBR2 composition at 200 MPa.
The highest silica concentrations were not found in the
higher pressure experiments. The experimental melts obtained from IBR2 can reproduce nearly all bulk compositions as a result of crystallization from 0 to 40 wt %.
The comparison of the compositions of experimental
glasses (Fig. 10; see also Supplementary File SM-1) and
of the average natural IBR2 glass indicates that the natural glasses are best reproduced in experiments containing 15 wt % crystals, which is consistent with the
crystallinity determined optically in the natural sample
(Fig. 1).
Figure 9 shows that the general trends observed in the
glasses of the Cougar Point Tuff system can be reproduced
experimentally using the BJR composition. However, the
composition of the natural sample IBR2 has higher TiO2,
FeO, MgO and CaO than the natural sample BJR
(see starting compositions in Figs 9 and 10), which strongly
suggests that the pre-eruptive temperature of the IBR2
composition was higher than that of BJR. Thus, in terms
of temperature and H2Omelt, BJR may be a differentiated
product of a magma with a composition similar to IBR2.
In this hypothesis, the temperature for the formation of
the bulk composition BJR must have been at least 9008C
(from experiments with IBR2, glasses with CaO and
TiO2 contents similar to the bulk composition BJR have
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ALMEEV et al.
SILICIC MAGMAS OF YELLOWSTONE HOTSPOT
been obtained only above 9008C). However, it is emphasized that the general temporal trend in the Bruneau^
Jarbidge rhyolites indicates that the most mafic rhyolites
erupted later than the evolved rhyolites. Thus, although
our experiments can be used to trace differentiation processes, a simple fractionation model to explain the chemical variability observed in the Bruneau^Jarbidge rhyolites
is not sufficient. Based on Sr^Nd isotope constraints, progressive addition of basaltic components (magma mixing)
to explain the general evolution from differentiated to
more mafic Fe^Mg-rich rhyolites can be ruled out
(Leeman et al., 2008). A possible explanation may be
related to the evolution of the composition of the crustal
source material that may become more refractory over
time, as a result of continuing melting, or a continuous increase of temperature with time in the protolith.
Although the natural glass composition of BJR can be
produced by differentiation from the bulk IBR2 starting
composition at 200 MPa, it is worth noting that the natural glass composition of BJR could not be reproduced
with the bulk composition BJR. For example, the Na2O,
CaO and Al2O3 concentrations of the natural BJR glasses
only partly overlap with the compositions obtained experimentally for the BJR composition (Fig. 9). As mentioned above, the low Al2O3 concentration (11·8 wt %)
(along with the low Na2O) of the starting BJR composition (Fig. 10) probably explains the absence of plagioclase
in most experiments and the discrepancy between the
compositional trends of the experimental and natural
glasses in Fig. 9. Thus, the bulk composition BJR may represent a sample from which part of the plagioclase has
fractionated. Although such a process may be difficult
considering the high viscosity of water-poor melts, fractionation may occur at a magmatic stage or during explosive eruption, transport and deposition (liquid^crystal
segregation leading to variable proportions of feldspars
to glass in different zones of the deposits). Alternatively,
if the phenocrysts and the silicate melts of the BJR
sample come from different regions of the magma reservoir(s), the proportion of plagioclase incorporated in the
silicate melt may have been lower than that of the other
phases (e.g. Al-poor mafic phases). Another possible alternative is that the bulk composition BJR may derive from
the partial melting of a protolith with a slightly different
composition than that of most rhyolitic samples. Finally,
it can also not be excluded that the chemical composition
of the BJR sample that was used as a starting material
was slightly modified by post-eruptive processes such as
increase of K2O at the expense of Na2O, although this is
thought to be unlikely in the dense, unhydrated vitrophyre sample. Alkali exchange or dilution most often
occurs in volcanic glass that has undergone hydration
through prolonged interaction with meteoric waters
(Cerling et al., 1985).
Potential of Qtz, Ab and Or proportions in
glasses to constrain pressure and melt
water content
The normative Qtz, Ab and Or proportions of the natural
glasses may be useful in constraining, at least qualitatively,
the pressure and water activity prevailing during magma
evolution. Considering that CaO contents are extremely
low in the rhyolites and that several natural glasses in
the Bruneau^Jarbidge rhyolites are in equilibrium with
quartz, sanidine and plagioclase (Ab-rich), their composition should be close to the minimum or eutectic point of
the haplogranitic Qtz^Ab^Or system. The glass compositions from experimental runs of BJR that coexist with
three tectosilicates are shifted towards the Qtz^Or sideline
of the ternary diagram when compared with the compositions of minima and eutectic points determined experimentally in the haplogranitic system at water-saturated
conditions. The same observation can be made with experimental glasses from IBR2 coexisting with quartz, sanidine and plagioclase (black symbols in Figs 11 and 12).
Strictly, Ca- and Fe-bearing rhyolitic compositions should
be projected into the synthetic Qtz^Ab^Or system. A correction for the An content was applied to account for
these additional elements following the procedure proposed by Blundy & Cashman (2001). This correction tends
to increase the Ab content at the expense of Qtz and Or
when compared with a projection using only the Qtz, Ab
and Or normative proportions (neglecting the An component). Thus, the shift of the glass compositions from experimental runs of BJR that coexist with three tectosilicates
away from the Ab end-member towards the Qtz^Or sideline cannot be due to the correction for the An content following Blundy & Cashman (2001). Based on experimental
data obtained in water-undersaturated haplogranitic systems, this observation is best explained by the strongly
reduced water-activity prevailing in the natural and experimental systems, resulting in a shift of the minimum
and eutectic points towards more Or-rich compositions
(Pichavant, 1987; Holtz et al., 1992; see evolution of the minimum composition with decreasing water activity at 200
and 500 MPa in Figs 11 and 12). It is worth noting that the
corrected Qtz^Ab^Or contents of the IBR2 experimental
glasses coexisting with quartz, sanidine and plagioclase
(black symbols in Fig. 12) plot along the trend connecting
the water-saturated and water-undersaturated minimum
and eutectic compositions at a given pressure (for
200 MPa as well as for 500 MPa). This trend indicates
that decreasing water activity results in a decrease of the
Ab/Or ratio at constant Qtz content and that decreasing
pressure results in an increase of the Qtz component in
the melt. The Or content of the IBR2 experimental glasses
is higher than that found in the haplogranitic system for
water-undersaturated conditions by Holtz et al. (1992),
which is consistent with the lower water activity of our
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JOURNAL OF PETROLOGY
VOLUME 53
experiments when compared with their data. Our data obtained for the BJR glasses confirm the observation above.
However, four glasses produced at 825 and 8758C have
higher Qtz contents than the haplogranitic minimum compositions at 200 MPa. The reason for this is unclear but
may reflect difficulties in the analysis of glasses in run
products with high crystallinity.
The good agreement between the Qtz^Ab^Or content
of experimental glasses coexisting with quartz, sanidine
and plagioclase synthesized from natural Ca-bearing rhyolitic compositions (this study) and the Qtz^Ab^Or content
of the minimum and eutectic points in the Ca-free haplogranitic system indicates that the Qtz content (corrected
for An content) may be used to constrain pressure (depth
of magma storage). Using this approach with the composition of glasses from the Bruneau^Jarbidge eruptive
center, it can be noted that the Qtz^Ab^Or contents of natural rhyolitic glasses in the IBR2 sample plot along a
trend with approximately constant Qtz content between
the 500 and the 200 MPa water-undersaturated cotectic
lines (crosses and grey field in Fig. 12). The interpolated
data suggest a pressure of 300^400 MPa (9^12 km),
which is consistent with pressures (5500 MPa) deduced
from the phase relationships (see above). Lower normative
Qtz in the CPT fallout glasses is more consistent with
slightly lower pressures (200^300 MPa; 10 km) for these
eruptive units (Fig. 11). This pressure range is consistent
with previous estimates (e.g. Leeman et al., 2008).
NUMBER 9
SEPTEMBER 2012
previous researchers (Honjo et al., 1992; Perkins & Nash,
2002; Cathey & Nash, 2004, 2009; Ellis et al., 2010).
The depth of magma storage prior to eruption remains
difficult to constrain. Preliminary data based on the stability of tectosilicate phases and on the glass compositions of
the natural samples indicate that pressure may have
varied in the range 400^200 MPa for the investigated compositions. In the absence of amphibole, pressure can best
be constrained from the stability fields of the tectosilicates.
However, the possible presence of xenocrysts (minerals
that were incorporated into the magma at higher crustal
levels) or possible segregation of phenocrysts from melt
during explosive eruptive processes makes it difficult to determine the depth of magma storage from phase equilibria
only. The determination of the composition of the glasses
quenched from silicate melts coexisting with two or three
tectosilicate phases may be an alternative to constrain
pressure in high-silica rhyolites, provided that the cotectic
and eutectic melt compositions obtained for natural rhyolitic systems at water-undersaturated conditions are used to
interpret the data.
AC K N O W L E D G E M E N T S
The original paper benefited from critical reading and
helpful comments by Jim Beard, Mike Carroll, Eric
Christiansen and Bill Leeman. We thank O. Dietrich for
preparing the thin sections, Wanja Dziony and Ju«rgen
Koepke for assistance in operating the electron microprobe, and Lars Kuschel, Sarah Cichy, Adrian Fiege and
Clemens Kirchner for technical assistance.
CONC LUSIONS
Our phase equilibria studies of the Bruneau^Jarbidge
rhyolites (less evolved IBR2, more evolved BJR) indicate
that the mineral assemblages observed in the natural samples can be reproduced experimentally only if the water
concentration in the melt is low and magmatic temperature is relatively high. At such conditions (low water activity), the redox conditions of our experiments reproduce
those determined from natural assemblages (QFM,
Cathey & Nash, 2004). Assuming temperatures of
900^9758C, the water content in the melts may be as low
as 1^0·6 wt % H2O, which confirms the hypothesis of the
nearly anhydrous character of the silicic magmatism of
the central Snake River Plain (see recent overviews by
Bonnichsen et al., 2008; Branney et al., 2008; Christiansen
& McCurry, 2008). By contrast, the crystallization of
amphibole in the Lava Creek Tuff at Yellowstone implies
higher magmatic water concentrations and lower temperatures. Amphibole (and biotite) crystallized in the BJR experiment at 8258C and 4·8% H2O, but biotite alone
crystallized at lower water contents than 2·9% H2O
in the IBR2 composition. The experimental data also
confirm the predicted high pre-eruptive temperatures
(900^10008C) obtained from mineral thermometry by
FU NDI NG
This study was funded by the German Research
Foundation (DFG, project Ho 1337/17 and Ho 1337/22).
B. Nash very much appreciates the award of a Mercator
Professorship and support from the German Research
Foundation during her sabbatical at the Leibniz University
of Hannover, as well as partial support from the National
Science Foundation, grant EAR0510327.
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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