Magma Recharge and Crystal Mush Rejuvenation

Transcription

JOURNAL OF PETROLOGY
VOLUME 50
NUMBER 11
PAGES 2095^2125
2009
doi:10.1093/petrology/egp070
Magma Recharge and Crystal Mush
Rejuvenation Associated with Early Post-collapse
Upper Basin Member Rhyolites, Yellowstone
Caldera, Wyoming
GUILLAUME GIRARD* AND JOHN STIX
EARTH AND PLANETARY SCIENCES, MCGILL UNIVERSITY, 3450 UNIVERSITY STREET, MONTREAL, QC, H3A 2A7,
CANADA
RECEIVED MARCH 10, 2009; ACCEPTED SEPTEMBER 23, 2009
ADVANCE ACCESS PUBLICATION NOVEMBER 3, 2009
The Upper Basin Member rhyolites are the oldest known postcollapse rhyolites of the third Yellowstone caldera. They were erupted
near the caldera’s resurgent domes between 516 7 and 473 9 ka
and at 257 13 ka. An unusual characteristic is their low d18O signature. Few data are available on their mineralogy and glass geochemistry, and this study fills an important gap in understanding
their petrogenesis.We report new mineralogical observations and plagioclase, whole-rock and glass compositional data. Based on our
observations, we describe a new lava flow for which we propose the
name East Biscuit Basin flow. This unit is a quartz- and sanidine-free low-silica rhyolite (71^72% SiO2 in the whole-rock) in
which the dominant mineral, plagioclase, comprises two populations:
(1) small fresh euhedral crystals of An20^48 (average of An31) composition, commonly part of aggregates with pyroxenes and Fe^Ti
oxides; (2) large sieve-textured isolated crystals, which are slightly
more sodic in composition (An19^34, average of An27). Plagioclase
compositions in most other Upper Basin Member rhyolites are similar. The range of compositions for trace elements such as Rb, Th,
Y and the rare earth elements is small (e.g. 158^189 ppm Rb, 20^
25 ppm Th, 52^63 ppm Y, 60^82 ppm La in the whole-rock), and
there is no systematic variation of these elements as a function of
SiO2 content or mineralogy. Certain trace element signatures and
ratios are specific to each of these rhyolites, allowing us to propose
that the Upper Basin Member rhyolites originate from six independent magma batches. The coexistence of the two types of plagioclase
and their progressive disappearance in the more evolved rhyolites
suggest the following petrogenetic model for each magma batch.
*Corresponding author. Telephone: þ1-514-398-5391.
Fax: þ1-514-398-4680. E-mail: [email protected]
A low-d18O rhyolitic protolith is heated by replenishing magmas,
which initiate melting, forming a crystal mush. Replenishment by
buoyant silicic magma may enhance melting and cause mixing with
the mush material. As a consequence, crystal-poor eruptible magma
batches are formed, which contain small, more calcic plagioclase
crystals (aggregates) formed during cooling and mixing of the
replenishing silicic melt, and larger, lower temperature crystals exhibiting dissolution features inherited from the protolith.
KEY WORDS: caldera volcanism; crystal transfer; melting; sieve
textures; silicic magma
I N T RO D U C T I O N
The period following cataclysmic eruptions and caldera
collapse in large rhyolitic systems is fundamental to our
understanding of magma chamber dynamics at these volcanic centres. Caldera collapse is commonly followed
by resurgence and post-caldera rhyolitic volcanism, which
is typically effusive (Smith & Bailey, 1968). Recent studies
at the Valles caldera, New Mexico, have revealed the very
short timescale (27 ka) in which caldera collapse, resurgence and post-caldera volcanism occur (Phillips et al.,
2007). Post-caldera eruptions are commonly perceived as
small volume, but at Long Valley caldera, California, the
Early Rhyolite that was emplaced in the 100 kyr period
The Author 2009. Published by Oxford University Press. All
rights reserved. For Permissions, please e-mail: journals.permissions@
oxfordjournals.org
VOLUME 50
following caldera collapse and eruption of the Bishop Tuff
totals 4100 km3 in volume (Hildreth, 2004). This period,
together with the late stages of large, caldera-forming
eruptions, is typically when the most compositionally
primitive and diverse silicic magma found in large silicic
systems is erupted, and compositions are commonly rhyodacitic. For example, at the Long Valley caldera a smallvolume rhyodacitic dome was erupted in the northeastern
part of the caldera soon after its collapse (Hildreth, 2004).
At the Valles caldera, a low-silica rhyolite lacking quartz
(the Redondo Creek lavas) was erupted immediately after
caldera collapse (Phillips et al., 2007). In New Zealand, at
the Taupo (Sutton et al., 2000) and Okataina volcanic centres (Smith et al., 2004; Shane et al., 2005), post-collapse
rhyodacitic eruptions are common, whereas the calderaforming eruptions emit mainly high-silica rhyolite.
It has long been debated whether these more primitive
magma types are the non-erupted remnants of a large
zoned magma chamber that was partially emptied during
the caldera-forming eruption, or new replenishing melts
that erupted without prolonged storage and differentiation.
Recent work on the oxygen isotope systematics of the
post-collapse rhyolites of the Yellowstone^Snake River
Plain volcanic province and the Oasis Mountain^Timber
Valley volcanic complex of Nevada led Bindeman &
Valley (2000, 2001) and Bindeman et al. (2006, 2007) to
introduce the concept of bulk melting of sub-solidus rhyolitic material inherited from previous eruptions as an alternative mechanism to explain the origin of such rhyolites.
This work focuses on the early post-collapse rhyolites of
the 640 ka Yellowstone caldera, Wyoming, also referred to
as the Upper Basin Member (UBM) (Christiansen, 2001).
These rhyolites were initially interpreted as remnants of
non-erupted magma from the chamber that produced the
Lava Creek Tuff (Hildreth et al., 1984). More recently,
Bindeman & Valley (2000, 2001) and Bindeman et al.
(2008) provided evidence to support bulk melting of precollapse rhyolites and subsided blocks of ignimbrite. This
work examines these UBM magmas, the relationships
between the various UBM units, and the role and mechanics of crystal entrainment. We propose that bulk melting
does not necessarily exclude other petrogenetic processes
such as fractional crystallization or magma mixing, and
that these processes may operate simultaneously with or
subsequent to bulk melting.
In this study, we report new whole-rock, glass, and plagioclase compositions for most of the known UBM rhyolites, to decipher whether the UBM rhyolites are
cogenetic or not. Spatio-temporal trace element distributions allow us to propose the existence of at least six independent magma batches. Our results generally confirm
that melting played a fundamental role in the generation
of these rhyolites, but they also provide evidence for
additional petrogenetic processes. In some rhyolites, a
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NOVEMBER 2009
combination of unusual mineral assemblages, coupled
with trace element heterogeneities in the glass, suggests
mixing with a small proportion of a more primitive silicic
magma. Furthermore, unusual trace element signatures in
certain rhyolites reflect heterogeneity of their source.
Based on the mineralogical and geochemical data presented, we provide evidence for a previously unrecognized
quartz- and sanidine-free low-silica rhyolite lava unit in
the Biscuit Basin area, which we name the East Biscuit
Basin flow. This flow contains plagioclase^pyroxene^Fe^
Ti-oxide microcrystalline aggregates with a bulk silica
content of 60% SiO2, as well as sieve-textured plagioclase. These features are also found in other UBM rhyolites, although they are less abundant.
E A R LY P O S T- C A L D E R A
VO L C A N I S M AT Y E L L O W S T O N E
After collapse of the third Yellowstone caldera and eruption of the 1000 km3 Lava Creek Tuff at 639 2 ka
(Lanphere et al., 2002), the Mallard Lake Resurgent
Dome formed in the southwestern part of the caldera, and
the Sour Creek Resurgent Dome formed in the northeastern part of the caldera (Christiansen, 2001) (Fig. 1). The
timing of these resurgence events is poorly constrained;
nevertheless, the earliest post-collapse rhyolites are found
near these two domes, and field relations suggest that they
were emplaced after resurgence (Christiansen, 2001). The
oldest known flow, the Biscuit Basin flow, has been dated
at 516 7 ka and lies astride the Mallard Lake Resurgent
Dome (Christiansen, 2001) (Fig. 1). A detailed d18O survey
by Bindeman & Valley (2000, 2001) redefined this unit as
a group of three flows: the South, Middle and North
Biscuit Basin flows (SBB, MBB and NBB, respectively),
with the SBB flow dated by the 39Ar/40Ar method at
255 11 ka (Bindeman et al., 2008). Here, we refer to these
rhyolites as the Biscuit Basin rhyolites. In the vicinity of
the Sour Creek Resurgent Dome, a sequence of two rhyolitic pyroclastic fallout units and two rhyolitic lava flows
was erupted between 486 42 and 473 9 ka (Gansecki
et al., 1996; Christiansen, 2001) (Fig. 1); we refer to these
as the Canyon rhyolites. Activity resumed near
the Mallard Lake Resurgent Dome with emplacement of
the Scaup Lake lava flow (SL) at 257 13 ka
(Christiansen et al., 2007) and the SBB flow at 255 11 ka
(Bindeman et al., 2008) (Fig. 1). Therefore, early postcaldera volcanism at Yellowstone appears to have occurred
relatively late after collapse compared with some other
silicic systems (e.g. Hildreth, 2004; Smith et al., 2005;
Phillips et al., 2007).
The UBM rhyolites are followed by the 600^900 km3
Central Plateau Member (CPM) rhyolites that were
emplaced in the Yellowstone caldera from 174 to 70 ka
(Christiansen, 2001; Christiansen et al., 2007) (Fig. 1).
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GIRARD & STIX
YELLOWSTONE UPPER BASIN MEMBER RHYOLITES
Fig. 1. (a) Location of Yellowstone caldera within the Yellowstone^Snake River Plain volcanic province (modified from Perkins & Nash, 2002).
Shaded areas represent the Yellowstone^Snake River Plain eruptive centres and their ages of activity. OH, Owyhee^Humboldt; BJ, BruneauJarbidge; TF, Twin Falls; P, Picabo; H, Heise; Y, Yellowstone. (b) Map of the Yellowstone caldera complex, showing the first and third calderas,
the resurgent domes and post-caldera volcanism (modified from Bindeman & Valley, 2001). (c) Map of the Lower Geyser Basin and Upper
Geyser Basin areas of the Yellowstone caldera, where the Biscuit Basin Rhyolites crop out, modified from US Geological Survey 1: 24 000 scale
topographic maps of Old Faithful, WY, and Madison Junction, WY, quadrants.
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VOLUME 50
These represent some of the largest rhyolitic effusive eruptions on Earth. Hence unlike other caldera systems
(Hildreth, 2004), the early UBM volcanism appears to represent a small proportion of the magma erupted after caldera collapse. However, the subsequent CPM rhyolites
have filled the caldera, so that volumes and numbers
of eruptive units for the UBM rhyolites may be
underestimated.
The UBM rhyolites have been recognized as the most
primitive rhyolites of the entire third caldera cycle at
Yellowstone (Hildreth et al., 1984; Christiansen, 2001). In
contrast to the pre-caldera lavas, the Lava Creek Tuff,
and the CPM rhyolites, plagioclase is the dominant mineral phase in the UBM rhyolites, whereas sanidine and
quartz are less abundant. Plagioclase ranges mainly from
An22 to An32 although values ranging from An18 to An43
have been reported (Hildreth et al., 1984). Clinopyroxene
is the dominant mafic phase and is relatively Mg-rich
(Wo38^40En27^35Fs25^35)
(Hildreth
et
al., 1984).
Orthopyroxene (En35^45) is also observed in the UBM
rhyolites whereas hydrous mafic minerals are not observed,
similar to most Yellowstone rhyolites (Hildreth et al., 1984).
Although less evolved, the UBM rhyolites are dominantly
high-silica rhyolites, although a few low-silica rhyolite
analyses also have been reported for the Biscuit Basin
rhyolites (Sturchio et al., 1986; Christiansen, 2001).
Eruption temperatures of 900^9108C have been determined by ilmenite^titanomagnetite geothermometry for
the 516 ka Biscuit Basin flow, 9238C for the 484 ka
Canyon flow and 850^8808C for the 257 ka Scaup Lake
flow (Hildreth et al., 1984). Zircon saturation temperatures
range from 819 to 8828C throughout the UBM rhyolites
(Bindeman & Valley, 2001).
The most unusual aspect of these rhyolites is their exceptionally low d18O values. Many of the UBM rhyolites have
values attaining 0^2ø d18O SMOW in quartz (Hildreth
et al., 1984). Similar depletions are found in glass, plagioclase, sanidine and zircon (Sturchio et al., 1986; Bindeman
& Valley, 2001). Although all post-caldera rhyolites from
the third Yellowstone caldera cycle have relatively low
d18O values (56·5ø in quartz), only the UBM rhyolites
have such low values. Zircon cores show higher values similar to the Lava Creek Tuff or older rhyolites, with U^Pb
ages on these zircons ranging between 42 Ma and 500
ka (Bindeman & Valley, 2001; Bindeman et al., 2001).
Hildreth et al. (1984) first interpreted these strong d18O
depletions as the result of contamination of the magma
chamber by meteoric water, whereas Bindeman & Valley
(2001) favoured bulk melting of hydrothermally altered
intra-caldera rhyolites. Elevated U, Th and Y concentrations in some zircons may indicate crystallization in a plutonic environment (as pegmatites or residual melts)
rather than a liquid magma chamber (Bindeman et al.,
2008).
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NOVEMBER 2009
S A M P L I N G A N D A N A LY T I C A L
M ET HODS
We collected samples at several localities from the Biscuit
Basin rhyolites near the Mallard Lake Resurgent Dome,
and the tuff of Sulphur Creek (TSC), Canyon flow (CF)
and Dunraven Road flow (DR) near the Sour Creek
Resurgent Dome as mapped by Christiansen (2001). For
the Biscuit Basin rhyolites, we sampled each of the North,
Middle and South flows as defined by Bindeman & Valley
(2001), as well as four localities east of the South Biscuit
Basin sample site (Table 1, Fig. 1). For these localities,
based on a mineralogy and geochemistry distinct from the
other Biscuit Basin flows, we define a new unit, the East
Biscuit Basin flow (or group of flows), hereafter referred
to as EBB. Samples of all units are typically glassy and perlitic with little or no spherulite development.
Crystal abundances were obtained by area measurements on thin-section scans. Whole-rock major and trace
element analyses were obtained by X-ray fluorescence
(XRF) on a Philips PW2440 4 kW spectrometer at
McGill University. Rare earth elements (REE) were
obtained by inductively coupled plasma mass spectrometry
(ICP-MS) using a PerkinElmer/SCIEX Elan 6100
DRCplus instrument at McGill University. An internal
standard, UTR-2, was used for both XRF and ICP-MS
analyses (analytical results are given in Electronic
Appendix 1, which is available for downloading at http://
www.petrology.oxfordjournals.org/). Major elements for
glass and plagioclase were obtained on a JEOL 8900 electron microprobe at McGill University using 15 kV accelerating voltage and 10 nA beam current. For glass analyses
we used a 10 mm defocused beam to minimize sodium
loss. Trace element analyses for glass were obtained by
laser-ablation (LA)-ICP-MS at McGill University using a
New-Wave UP-213 laser coupled to the ICP-MS system.
For LA-ICP-MS analyses, the standard used for calibration
was NIST-610, whereas NIST-612 was analyzed to monitor
accuracy and precision. Both standards were analyzed several times during the day, approximately once every hour.
The beam size was 60 mm for all analyses to ensure relatively constant count rates, and drilling was performed to
a depth of 60 mm. Raw isotope counts were converted to
elemental concentrations using Glitter software; our reference isotope of known concentration was 29Si rather than
the more commonly used 43Ca because of the low calcium
concentrations in the glasses. However, 29Si interferes with
the atmospheric polyatomic ion 14N15N, and a method
that minimizes atmospheric contamination was required.
The laser-ablation chamber can contain a maximum of
one thin section; hence changing thin sections caused
atmospheric contamination that altered the instrumental
accuracy and precision. Instead of thin sections, therefore,
we used polished mounts containing millimeter-sized
obsidian chips from 10^15 different samples including
2098
Map
abbreviation
sample
name
Qpub
BS88-05
2099
Qpub
Qpul
BN154A
BS132
SL32-05
Qpus
Qpud
Qpud
TSC217-2
DR71-6
DR237B
Dunraven Road flow
Dunraven Road flow
Tuff of Sulphur Creek
Canyon flow
Scaup Lake flow
South Biscuit Basin flow
North Biscuit Basin flow
Middle Biscuit Basin flow
East Biscuit Basin flow
Name
DR
DR
TSC
CF
SL
SBB
NBB
MBB
EBB
EBB
EBB
EBB
Abbreviation
Ar–Ar
486 42
486 42
473 9
484 15
257 13
255 11
n.d.
516 7
n.d.
n.d.
n.d.
n.d.
age (ka)
542103
542348
546753
539295
515025
510173
514125
513762
512216
512002
512526
511080
x
4956557
4954868
4956350
4952953
4921761
4925475
4932331
4929922
4924296
4924186
4924108
4925636
y
US UTM reference
NAD1927 Continental
Sampling location
Geographical
Fresh roadcut
vent area
Top of butte near mapped
50 cm above exposed base
North side of creek
North rim of canyon
Along trail at Mystic Falls
Top of butte
Top of butte
North side of butte
Top of butte
description
Crystal-poor obsidian
Crystal-poor obsidian
Vitrophyre pyroclastic fallout
Perlitic spherulite-rich obsidian
Perlitic obsidian, frothy flow carapace
Perlitic obsidian
Perlitic obsidian
Perlitic obsidian
Perlitic obsidian
Lithology
Previous sampling
8YC-477B1,3
YL96-92
YL96-22
BBF-12
6YC-1061
at same locality
Sources for ages: Gansecki et al. (1996), except for SBB (Bindeman et al., 2008) and SL (Christiansen et al., 2007, from A.T. Calvert, unpublished data). Names for
the Biscuit Basin flows as proposed by Bindeman & Valley (2001), except for East Biscuit Basin flow, this study. Map abbreviations from Christiansen (2001). Data
from previous sampling published as follows: 1Christiansen (2001); 2Bindeman & Valley (2001); 3Hildreth et al. (1984). n.d., not determined.
Qpuc
CF133B
Canyon rhyolites
Qpub
Qpub
BM7-3
Qpub
Qpub
BS87
BS89
Qpub
BS47
Biscuit Basin rhyolites
Unit
Group,
Table 1: Stratigraphy and published 39Ar/ 40Ar ages of the Upper Basin Member rhyolites, description of samples and sampling localities
GIRARD & STIX
VOLUME 50
standards. Obtaining 10 analyses for each sample on the
mount required about 1 day; thus the sample chamber
was kept closed during an entire analytical session. We
used the average SiO2 content in glass obtained by electron
microprobe as the SiO2 content of reference. For electron
microprobe analyses, variations in SiO2 content within a
single thin section were consistently within instrumental
uncertainty, demonstrating that the glass was homogeneous in its SiO2 content at the thin-section scale.
A few samples were not analyzed by electron microprobe. For these, we used the average SiO2 content of our
analyses for other units in the same stratigraphic group
for which SiO2 could be considered homogeneous and representative. Analyzed isotopes comprised 85Rb, 88Sr, 89Y,
90
Zr, 137Ba, 139La, 147Sm, 153Eu, 157Gd, 172Yb, 175Lu, 232Th,
and 238U.
R E S U LT S
Mineralogy
Our mineralogical observations and calculated crystal
abundances are reported in Table 2. In terms of mineralogy, the most primitive material comes from the East
Biscuit Basin flow (EBB, samples BS88-05 and BS89,
Table 1). This lava has 15% crystals and is dominated by
plagioclase whereas sanidine and quartz are absent.
Clinopyroxene is the second most abundant mineral. Such
a crystal assemblage is uncommon for a rhyolite with
71^72% SiO2 in the whole-rock and 76% SiO2 in the
glass (compositions normalized to 100% oxides).
Plagioclase exhibits two types: (1) 1mm long euhedral
crystals, commonly optically zoned, occurring either
singly (Fig. 2a) or intergrown with pyroxenes and Fe^Ti
oxides, forming microcrystalline aggregates (Fig. 2b); (2)
3^5 mm long euhedral sieved crystals typically occurring
as single grains. In these, the sieve pattern is variable,
with a majority of coarse sieve textures (Fig. 2c) and
fewer fine sieve textures (Fig. 2d). Rims on sieved crystals
are not systematically present; when observed, they are
only partly developed around the sieved core. The aggregates represent 30% of the crystallinity of the EBB flow
and are composed of 75% plagioclase, 20% pyroxenes
and 5% Fe^Ti oxides. They are typically 1mm in size
with the largest attaining 5 mm. The presence of glass in
the aggregates is not evident.
Middle Biscuit Basin flow (MBB) sample BM7-3
(Table 1) is somewhat less crystalline (12% crystals),
with 40% sanidine and 10% quartz, although plagioclase is still dominant at 50% (Fig. 3). Sieved plagioclase
is also abundant, but sieve textures are generally less developed than for the EBB flow. Plagioclase^pyroxene^oxide
aggregates are scarcer (20% of the crystallinity); in
these aggregates, mafic minerals are less abundant than
for the EBB samples. Texturally, quartz is typically
rounded with long glass re-entrants, appearing to be in
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NOVEMBER 2009
dissolution (Fig. 2e). North Biscuit Basin flow (NBB)
sample BN154A (Table 1) generally resembles MBB, but
here sanidine is the dominant mineral (450%), with the
largest crystals being sieved. Plagioclase sieve textures are
less pronounced. Aggregates are also scarcer and represent
only 10% of the crystallinity (Table 2). Quartz is texturally similar to that of the MBB flow. South Biscuit Basin
flow (SBB) sample BS132 (Table 1) resembles NBB in
terms of quartz, sanidine and plagioclase abundances and
textures. However, the presence of aggregates in BS132 is
unclear although it is pyroxene-rich (5%). Indeed, both
orthopyroxene and clinopyroxene occur here as larger, isolated, euhedral crystals. sometimes approaching 1mm in
size. This mineralogy is similar to that of the 257 ka
Scaup Lake flow (SL) sample SL32-05 (Fig. 3). Electron
microprobe analyses reveal that these clinopyroxenes are
compositionally similar to those of the aggregates in the
MBB flow, with Wo42^43En32Fs25^26 in SL, Wo38^42En28^35
Fs25^32 in SBB and Wo39^41En31^35Fs25^29 in MBB. These
values are consistent with those obtained by Hildreth
et al. (1984).
Among the Canyon rhyolites (Table 1), Canyon flow
(CF) sample CF133B is the most primitive in mineralogy,
consisting mainly of a plagioclase^quartz assemblage with
scarce sanidine. Sieve textures in plagioclase are conspicuously developed similar to the EBB flow, and smaller,
fresh euhedral plagioclase is also found, either isolated or
as part of plagioclase^pyroxene^oxide aggregates,
although these aggregates are rare (510% of the crystallinity). Quartz dissolution textures are also well developed.
The underlying tuff of Sulphur Creek (TSC) sample
TSC217-2 has a mineralogy similar to CF, except for a
lower crystallinity. Dunraven Road flow (DR) samples
DR71-6 and DR237B are distinct from the rest of the
UBM samples in terms of their nearly aphyric nature
(1^3% crystals), although plagioclase is still dominant
(Fig. 3). In general, too few crystals are present in the DR
samples for accurate crystal counting.
Whole-rock geochemistry
Whole-rock major and trace element compositions are presented in Table 2. The data generally agree with those of
Hildreth et al. (1984, 1991) and Christiansen (2001). When
normalized to 100% oxides, all samples are high-silica
rhyolites, except for the East Biscuit Basin (EBB) samples
BS47, BS88-05 and BS89 (71·5^72% SiO2). The Middle
Biscuit Basin (MBB) and North Biscuit Basin (NBB) samples have 76·1% SiO2, whereas the mineralogically
more evolved South Biscuit Basin and Scaup Lake samples
have 74·8^75·2% SiO2. No low-silica rhyolite is found in
the Canyon rhyolites. There, the most primitive material
is from the Canyon flow (CF) with 75·8% SiO2. The
tuff of Sulphur Creek (TSC) is more silicic at 76·8%
SiO2, and the Dunraven Road flow (DR) is the most silicic
UBM rhyolite at 76·8^77·0% SiO2.
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GIRARD & STIX
Table 2: Whole-rock XRF and ICP-MS analyses and mineralogy
Sample:
BS47
BS88-05
BS89
BM7-3
BN154A
BS132
SL32-05
CF133B
TSC217-2
DR71-6
DR237B
Unit:
EBB
EBB
EBB
MBB
NBB
SBB
SL
CF
TSC
DR
DR
XRF results
SiO2
TiO2
70·12
0·480
69·63
0·475
70·13
0·488
73·57
0·263
73·81
0·245
72·46
0·269
73·26
0·259
74·18
0·288
74·39
0·342
76·74
0·127
76·53
0·125
Al2O3
13·21
13·33
13·08
12·07
11·96
12·72
12·51
12·56
12·58
12·20
12·26
FeO
3·44
3·48
3·47
1·59
1·62
1·75
1·69
1·21
1·47
1·27
1·24
MnO
0·078
0·076
0·074
0·041
0·040
0·043
0·043
0·033
0·042
0·036
0·034
MgO
0·47
0·47
0·47
0·17
0·20
0·31
0·23
0·10
0·12
0·10
0·11
CaO
1·65
1·58
1·60
0·72
0·68
0·91
0·84
0·50
0·92
0·50
0·51
Na2O
3·58
3·67
3·47
3·07
3·08
3·28
3·30
2·74
3·44
3·52
3·55
5·20
K2O
4·51
4·46
4·60
5·18
5·24
4·98
5·18
4·99
4·84
5·21
P2O5
0·110
0·108
0·110
0·032
0·032
0·035
0·036
0·037
0·036
0·018
BaO (ppm)
LOI
Total
Ga
1110
1075
1105
925
831
1252
1255
910
1165
709
2·31
2·11
2·45
3·32
3·24
3·01
2·51
1·70
3·35
0·19
0·23
100·07
99·50
100·06
100·12
100·23
99·89
99·98
100·08
99·99
99·98
99·88
21·6
21·1
21·4
18·2
115·8
111·5
111·4
49·2
46·3
59·8
57·9
82·5
45·4
17·8
62·5
60·2
60·0
60·1
60·3
56·1
57·7
61·7
60·1
52·0
Nb
Ce
41·3
162
41·2
146
41·6
153
45·4
146
280
44·6
151
42·3
161
262
41·6
146
389
46·9
147
175
18·4
Sr
280
164
19·0
Y
292
171
19·5
160
426
171
19·1
158
415
189
19·6
157
413
187
18·6
Rb
Zr
0·017
738
340
48·8
156
168
205
33·4
127
18·3
169
17·1
52·3
205
33·4
112
Pb
26·3
26·3
25·7
30·6
30·8
25·5
26·4
28·9
27·7
30·5
30·9
Th
21·8
21·6
21·4
24·7
24·5
23·4
23·1
22·3
24·8
19·8
19·6
U
5·1
5·6
5·1
5·9
5·8
5·2
5·3
5·5
5·4
4·0
4·2
ICP-MS results
Ba
n.a.
n.a.
La
n.a.
n.a.
936
73·7
73·4
75·7
80·7
81·8
75·8
72·5
n.a.
60·0
Ce
n.a.
n.a.
139·8
139·9
144·3
151·0
152·6
144·6
141·0
n.a.
114·0
Pr
n.a.
n.a.
16·0
15·7
16·1
16·9
16·9
16·5
16·3
n.a.
12·9
Nd
n.a.
n.a.
59·9
56·6
58·1
60·5
61·1
60·7
60·3
n.a.
47·1
Sm
n.a.
n.a.
11·9
11·1
11·3
11·2
11·3
11·9
12·0
n.a.
9·37
Eu
n.a.
n.a.
n.a.
0·62
Gd
n.a.
n.a.
n.a.
8·40
Tb
n.a.
n.a.
n.a.
1·43
Dy
n.a.
n.a.
n.a.
8·88
Ho
n.a.
n.a.
2·14
2·07
2·11
1·98
2·01
2·16
2·15
n.a.
1·75
Er
n.a.
n.a.
6·16
5·97
6·19
5·68
5·83
6·21
6·36
n.a.
4·97
Tm
n.a.
n.a.
0·89
0·90
0·91
0·82
0·85
0·92
0·91
n.a.
0·73
Yb
n.a.
n.a.
5·88
5·81
6·03
5·53
5·62
5·96
6·09
n.a.
4·73
Lu
n.a.
n.a.
0·84
0·83
0·84
0·79
0·79
0·85
0·87
n.a.
0·67
n.d.
n.d.
n.d.
2·01
10·7
1·77
11·0
800
1·16
9·88
1·72
10·4
758
1047
1·12
10·1
1·70
10·7
1038
1·50
1·49
9·70
9·75
1·61
1·64
9·89
10·1
1002
1·85
10·6
1·75
10·9
789
1·32
10·3
1·78
10·9
n.a.
616
Mineralogy
Crystallinity (vol. %)
15
12
14
12
14
18
5
11%
2%
2%
4%
5%
3%
—
—
3%
2%
2%
2%
2%
3%
1%
1%
(unsieved)
60%
24%
18%
10%
33%
31%
24%
(sieved)
27%
20%
14%
14%
4%
41%
31%
n.d.
—
39%
54%
52%
37%
5%
12%
25%
—
12%
11%
10%
18%
17%
32%
—
35%
20%
9%
?
?
13%
—
—
Pyroxenes
Fe–Ti oxides
Plagioclase
Sanidine
Quartz
Aggregates
1
75%
n.d.
Crystals present in aggregates are counted in the total abundance values for each mineral. Aggregates comprise 80%
unsieved plagioclase, 18% pyroxene and 3% oxides in EBB; 50% of each of these phases are part of aggregates.
Plagioclase abundance increases slightly at the expense of pyroxene in aggregates of MBB and NBB. For pyroxenes, no
detailed count of clinopyroxene vs orthopyroxene was obtained, although both are observed in all pyroxene-bearing
rhyolites, with clinopyroxene dominant. For SBB and SL, presence of aggregates is uncertain, as polycrystalline assemblages essentially consist of few small pyroxene and oxide inclusions in larger unsieved plagioclase. n.a., not analyzed;
n.d., not determined.
2101
VOLUME 50
NUMBER 11
NOVEMBER 2009
Fig. 2. Photomicrographs of typical Upper Basin Member crystals. (a) Zoned, unsieved, euhedral plagioclase, occurring as single grains (view
under crossed polars); (b) unsieved plagioclase (PL), occurring in aggregates with pyroxenes (PX) and Fe^Ti oxides (OX); (c) coarsely
sieved plagioclase; (d) less commonly occurring finely sieved plagioclase; (e) quartz exhibiting rounding and long glass re-entrants. All
images taken from sample BS89, East Biscuit Basin flow except (e) taken from sample CF133B, Canyon flow.
2102
GIRARD & STIX
20
18
16
modal %
14
12
10
8
6
4
2
0
EBB
MBB
NBB
SBB
SL
CF
TSC
Sieved plagioclase
Sanidine
Pyroxene
Unsieved plagioclase
Quartz
Fe-Ti oxides
DR
Fig. 3. Mineralogy of samples of the Upper Basin Member rhyolites. (For this and following figures, see Table 1 for abbreviations.)
All samples show a negative correlation between SiO2
and TiO2, Al2O3, FeO, MgO, CaO and Sr (Figs 4 and
5). In contrast, the alkalis, including Na2O, K2O and Rb,
are generally invariant (Figs 4 and 5). Most trace elements
do not show a simple univariant evolutionary trend. For
example, Ba in the primitive EBB samples is 1000 ppm,
comparable with CF (960^1080 ppm) but slightly lower
than in the SBB (1100^1120 ppm) and SL (1120 ppm)
samples (Fig. 5). All the UBM rhyolites have a well-defined
negative Eu anomaly, although it is relatively small
(Fig. 6); its presence suggests that all UBM magmas,
including the primitive EBB, have undergone some feldspar fractionation. Most analyzed trace elements lack any
clear variation with mineralogy and SiO2 content, showing small ranges in composition. OnlyTh increases slightly
with SiO2, from 21^22 ppm in EBB to 25^26 ppm in
MBB, NBB and TSC, although it is significantly lower in
DR (19·5^20 ppm) (Fig. 5).
No significant changes in REE abundances are observed
as a function of the degree of differentiation. For example,
the Ce/Yb ratio shows a limited range between 23 and
24·5 (Fig. 5). Only SBB and SL have a higher Ce/Yb ratio
of 27^28, resulting from a combination of slightly lower
concentrations in middle REE (MREE) and heavy REE
(HREE) and slightly higher concentrations in light REE
(LREE) (Fig. 6b, Table 2). Although DR is rich in SiO2, it
is depleted in all the REE and other trace elements
(except Rb) compared with other UBM rhyolites, in
addition to major oxides such as TiO2, Al2O3, FeO, MgO
and CaO (Figs 4^6). This observed depletion is generally
more pronounced for the elements compatible in feldspar
(Sr, Ba and Eu). This relationship is probably fundamentally important for the origin of this chemically unusual
lava flow, which has the highest SiO2, a relatively primitive
plagioclase-dominated mineralogy, and a depleted trace
element signature.
Glass geochemistry
Results for glass geochemistry are presented in Table 3. For
each element and sample, we provide the number of analyses, the average concentration, per cent relative standard
deviation, and minimum and maximum concentrations
measured. This approach allows us to document compositionally heterogeneous glass. Heterogeneity can be
revealed by the presence of bimodal compositions within
a sample (Cathey & Nash, 2004; Smith et al., 2004, 2005),
or by compositional variability for certain elements that
greatly exceeds instrumental precision. An average of
10 analyses was performed on each sample whenever
possible. Our criteria for determination of heterogeneity
are detailed in Electronic Appendix 2 (available at http://
www.petrology.oxfordjournals.org/).
Glass in all samples is high-silica rhyolite. On an anhydrous basis, SiO2 is 76^76·5% in the East Biscuit Basin
(EBB) low-silica rhyolites, 77·1% in the Middle
Biscuit Basin (MBB) and North Biscuit Basin (NBB), and
2103
0.7
NUMBER 11
2.5
(a)
0.6
NOVEMBER 2009
(b)
2
0.5
CaO (wt%)
T i O 2 (wt %)
VOLUME 50
0.4
0.3
0.2
1.5
1
0.5
0.1
0
0
71
73
74
75
76
SiO 2 (wt%)
77
78
71
(c)
Na2O (wt%)
4
3
2
1
73
74
75
76
SiO 2 (wt%)
77
78
72
73
74
75
76
SiO 2 (wt%)
77
78
72
73
74
75
76
SiO 2 (wt%)
77
78
(d)
4
3.5
3
2.5
0
2
71
0.8
72
73
74
75
76
SiO 2 (wt%)
77
78
71
6
(e)
5.5
0.6
K 2O (wt%)
MgO (wt%)
72
4.5
5
FeO (wt %)
72
0.4
(f)
5
4.5
4
0.2
3.5
0
3
71
72
73
74
75
76
SiO 2 (wt%)
77
Biscuit Basin Rhyolites
Canyon Rhyolites
78
71
EBB
MBB
NBB
CF
TSC
DR
SBB
SL
Fig. 4. Major element whole-rock compositions of the Upper Basin Member rhyolites, normalized to 100% oxides, compiled from our data
(Table 2), Hildreth et al. (1984, 1991), Bindeman & Valley (2001) and Christiansen (2001).
77·5^77·7% in South Biscuit Basin (SBB) and Scaup Lake
(SL) flows. Glass in the Canyon rhyolites (Table 1) has
76·7^76·9% SiO2 for the Canyon flow (CF) and tuff of
Sulphur Creek (TSC), and 77·3^77·5% SiO2 for the
Dunraven Road flow (DR). SBB and SL glasses are also
characterized by very low FeO (0·38%), whereas FeO
ranges between 1·0 and 1·7% in all other UBM rhyolites
glass.
Concentrations of trace elements are generally lower in
the glass than the whole-rock, with the notable exception
of Rb and, to a lesser extent, U and Th. Th is similar in
glass and whole-rock for the Biscuit Basin rhyolites (somewhat higher in the glass in the EBB flow) and lower in the
glass in the Canyon rhyolites, especially for TSC and DR.
Trace element depletions in the glass of the MBB flow are
generally less pronounced for Th and REE (except Eu),
being about equal in glass and whole-rock. Such a trace
element distribution pattern requires the presence of accessory mineral phases in which the analyzed elements are
compatible.
2104
GIRARD & STIX
150
250
(a)
90
60
150
100
30
50
0
71
72
1500
74
75
SiO2 (wt%)
76
77
0
71
78
600
500
Zr (ppm)
Ba (ppm)
73
(c)
1200
900
600
300
73
74
75
SiO2 (wt%)
76
77
78
73
74
75
SiO2 (wt%)
76
77
78
73
74
75
SiO2 (wt%)
76
77
78
73
74
75
SiO2 (wt%)
76
77
78
(d)
300
200
100
71
160
72
73
74
75
SiO2 (wt%)
76
77
0
71
78
72
28
(e)
150
140
Ce / Yb
Ce (ppm)
72
400
0
130
120
(f)
26
24
22
110
100
71
72
73
74
75
SiO2 (wt%)
76
77
78
20
71
72
7
27
U (ppm)
(g)
25
Th (ppm)
(b)
200
Rb (ppm)
120
Sr (ppm)
23
21
19
6
(h)
5
4
17
3
15
71
72
73
74
75
SiO2 (wt%)
76
Canyon Rhyolites
77
78
71
72
EBB
MBB
NBB
CF
TSC
DR
SBB
SL
Fig. 5. Trace element whole-rock compositions of the Upper Basin Member rhyolites, normalized to 100% oxides, plotted as a function of SiO2,
which reflects the degree of evolution, compiled from our data (Table 2), Hildreth et al. (1984, 1991), Bindeman & Valley (2001) and
Christiansen (2001).
2105
VOLUME 50
NUMBER 11
NOVEMBER 2009
1000
EBB
MBB
NBB
Rock / Chondrite
(a)
100
10
1000
Rock / Chondrite
(b)
EBB
SBB
SL
100
10
1000
EBB
TSC
CF
DR
Rock / Chondrite
(c)
100
10
Rock / "primitive" EBB
1.4
1.2
(d)
1
0.8
0.6
0.4
EBB
CF
SL
0.2
TSC
DR
NBB
0
La Ce Pr Nd Sm Eu Gd Tb Dy Ho
Er Tm Yb Lu
Fig. 6. Whole-rock REE plots for the Upper Basin Member rhyolites, normalized to chondrites (Sun & McDonough, 1989) (a^c), and to EBB
sample BS89, the most primitive Upper Basin Member rhyolite (d).
2106
GIRARD & STIX
Table 3: Glass electron microprobe and LA-ICP-MS analyses
Sample: BS47
BS88-05
BS89
BS87
Unit:
EBB
EBB
EBB
EBB
BM7-3
BN154A
BS132
SL32-05
MBB
NBB
SBB
SL
12
15
6
23
Electron microprobe analyses
11
n:
SiO2
TiO2
Al2O3
n.a.
73·97
n.a.
72·40
n.a.
0·24
n.a.
n.a.
1·22%
75·07
10%
72·55
71·23
0·24
1·14%
10%
0·31
0·21
0·28
11·81
1·3%
11·47
1·8%
12·01
11·13
9·6%
1·63
5·7%
1·20
1·56
1·46
1·73
MnO
n.a.
MgO
n.a.
0·05
CaO
n.a.
0·52
0·44
0·63
0·41
0·52
Na2O
n.a.
3·44
8·4%
3·19
2·5%
2·69
3·90
3·05
3·36
5·3%
5·58
1·3%
0
0·017
K2O
n.a.
5·19
P2O5
n.a.
0·02
Cl
n.a.
0·06
Total
n.a.
96·69
4·60
0·005
0·046
94·78
63%
0·036
43%
0·070
11%
5·44
55%
0·031
11%
0·064
1·35%
98·39
0·04
0·016
0·07
0·055
0·47
5·51
0·02
0·003
0·05
0·043
95·30
93·60
73·60
72·21
n.a.
0·20
n.a.
11·93
1·39
0·02
n.a.
74·47
0·22
11·52
FeO
10
39%
n.a.
0·14
1·82%
75·42
23%
73·97
73·33
0·15
0·23
0·09
0·18
0·10
11·35
1·1%
11·16
0·9%
11·63
1·7%
11·53
11·60
11·02
11·33
11·43
11·87
1·26
3·2%
0·36
16·0%
0·36
1·20
1·36
0·27
0·43
0·23
0·066
0·007
0·073
0
0·048
0
0·67
69%
0·054
0·39
0·47
0·43
0·48
0·35
0·50
0·32
0·51
3·26
5·3%
3·13
7·2%
3·27
4·4%
3·22
4·6%
2·89
3·45
2·55
3·33
3·13
3·49
2·92
3·53
n.a.
5·31
2·7%
5·20
3·8%
5·44
1·6%
5·50
2·9%
5·07
5·54
4·97
5·58
5·34
n.a.
0·01
107%
0·01
107%
0·01
0
n.a.
0·05
n.a.
95·47
0·044
97·98
93·96
0·022
8%
0·057
0·99%
97·48
0·049
0·45
0
0·06
0·055
94·12
93·25
0·083
5%
0·027
7%
0·069
0·61%
95·22
0·02
0
0·44
0
0·07
0·055
95·04
93·36
82%
0·02
n.a.
0·088
5%
0·06
15%
86%
0·44
0·066
9%
0·02
1·0%
29·8%
3·2%
1·21
58%
0·21
11·82
1·14
0·04
19%
11·26
1·10
47%
0·40%
74·45
n.a.
0·057
1·16%
72·44
0·08
0·047
11%
16%
73·72
0·11
0·009
5·69
62%
0·18
0·53%
73·38
n.a.
0·085
7%
72·06
0·32
0·04
0·061
16%
19%
72·58
0·16
11·20
n.a.
1·03%
75·13
0·037
14%
5·60
81%
0·018
19%
0·090
1·93%
97·40
0·01
0
0·37
115%
0·035
10%
5·17
5·81
0·00
171%
0
0·07
0·050
95·17
94·62
0·029
9%
0·073
0·36%
95·77
LA-ICP-MS analyses
n:
Rb
6
202
200·2
Sr
28·2
25·4
Y
47·2
43·8
Zr
272
255·5
Ba
765
736·7
La
68·6
65·4
Sm
Eu
Gd
Yb
Lu
Th
0·75% 202
204·7
5·7%
29·4
4·6%
49·7
199·8
27·3
25·2
49·1
46·5
3·61% 283
281·8
2·5%
780·6
2·9%
70·7
270·3
763
738·5
70·9
67·9
9·34
5·2%
9·76
8·38
9·75
9·07
0·93
7·5%
0·87
0·84
1·02
8·02
4·2%
7·66
5·04
12
1·06% 222
205·2
6·6%
30·1
4·3%
51·7
210·2
13·6
9·28
45·6
44·6
4·03% 260
296·7
3·0%
791·5
3·7%
74·0
588
436·1
66·8
65·6
3·07% 250
229·9
238·4
10
2·25% 219
257·6
215·5
27·0%
16·0
23·1%
12·3
5·43
12·2
14·7
16·7
11·7
1·3%
46·3
47·3
45·4
0·90% 272
263·7
12·2%
693·1
1·2%
67·8
264·7
417
299·9
69·3
64·5
3·9%
51·9
53·4
46·0
1·96% 213
280·8
23·0%
608·5
3·0%
72·6
359
326·2
74·8
65·5
8·79
11·07
8·94
5·7%
0·73
12·3%
0·57
14·0%
0·66
0·81
0·94
0·52
0·83
0·45
0·71
8·34
5·9%
7·82
4·9%
8·05
7·0%
8·64
7·54
8·85
7·27
8·76
6·84
6·9%
5·20
3·8%
4·74
5·7%
4·90
4·52
5·39
4·96
5·49
4·26
5·09
0·70
2·9%
0·72
4·2%
0·66
0·68
0·74
0·67
0·75
0·59
23·3
22·2
3·8%
24·3
21·3
20·6
6·9%
56·6
46·4
44·7
6·56% 173
225·3
4·6%
375·1
6·4%
79·2
230
220·8
67·0
65·4
0·86% 203
220·6
2·6%
12·8
2·1%
47·5
199·8
17·2
17·0
47·4
47·0
2·10% 147
178·3
2·0%
238·0
2·3%
69·9
146·0
368
364·3
68·1
67·4
10
1·24% 204
205·5
0·9%
17·4
0·8%
47·9
199·7
16·8
15·7
46·4
45·7
0·60% 144
148·0
1·1%
373·6
1·0%
68·8
140·3
369
354·5
66·4
65·2
2·12%
212·6
6·5%
19·7
0·9%
47·1
1·22%
146·2
1·9%
382·3
1·0%
67·2
5·5%
9·29
6·4%
9·06
8·46
9·74
8·43
9·70
8·42
9·63
7·6%
0·47
10·6%
0·79
13·9%
0·72
9·7%
0·57
0·72
0·39
0·52
0·65
0·91
0·62
0·80
8·77
6·6%
7·82
5·4%
7·75
2·6%
7·64
5·1%
8·86
7·96
9·73
7·30
8·57
7·51
7·94
6·95
8·19
8·4%
5·68
7·9%
4·99
5·0%
5·04
2·8%
4·94
2·8%
4·38
5·51
5·01
6·24
4·56
5·36
4·85
5·16
4·79
5·27
6·1%
0·69
4·3%
0·78
9·0%
0·66
3·0%
0·70
4·3%
0·67
4·5%
0·70
0·63
0·71
0·68
0·86
0·62
0·69
0·68
0·74
0·62
21·7
22·3
21·2
1·9%
22·7
10·41
25·8
22·8
6·3%
166·6
4
9·23
1·5%
6·6%
184·9
9·64
23·3
214·8
7·7
9·77
4·4%
1·19% 217
223·3
20·4
3·2%
10·47
10
23·5%
8·78
22·4
5·9%
256·5
10
9·20
20·8
U
6
11·21
6·4%
27·6
22·9
21·6
3·3%
24·3
22·9
22·8
0·9%
23·2
22·1
21·5
4·4%
0·72
1·9%
22·7
5·81
1·9%
5·98
1·8%
5·46
1·6%
5·79
1·9%
6·87
1·5%
6·71
1·5%
6·04
1·5%
5·99
2·8%
5·70
5·96
5·83
6·13
5·33
5·65
5·61
5·91
6·68
7·00
6·55
6·90
5·91
6·11
5·72
6·21
(continued)
2107
VOLUME 50
NUMBER 11
NOVEMBER 2009
Table 3: Continued
Sample:
Unit:
CF133B
CF
Electron microprobe analyses
n:
16
TSC217-2
TSC
DR71-6
DR
DR237B
DR
21
8
8
SiO2
76·10
72·83
TiO2
0·22
0·15
Al2O3
11·80
11·44
1·3%
11·95
11·41
11·33
0·5%
11·53
11·72
11·60
0·7%
11·83
11·80
11·67
0·7%
11·97
FeO
1·46
1·34
3·5%
1·54
1·36
1·14
5·5%
1·44
0·99
0·93
5·8%
1·11
1·08
1·01
5·3%
1·18
MnO
0·05
0·011
49%
0·074
0·04
0·001
47%
0·080
0·03
0
66%
0·049
0·02
0
81%
0·042
MgO
0·08
0·060
18%
0·109
0·07
0·043
21%
0·096
0·04
0·016
29%
0·052
0·06
0·040
17%
0·069
CaO
0·49
0·46
5%
0·54
0·42
0·34
12%
0·51
0·38
0·34
7%
0·42
0·42
0·37
7%
0·47
Na2O
3·49
3·31
2·3%
3·60
3·11
2·99
2·1%
3·21
3·68
3·62
0·9%
3·72
3·68
3·62
1·1%
3·75
K2O
5·25
5·14
1·4%
5·40
5·81
5·62
1·2%
5·91
5·17
5·12
0·7%
5·20
5·20
5·14
0·5%
5·24
P2O5
0·01
0
133%
0·043
0·01
0
99%
0·038
0·01
0
86%
0·022
0·01
0
140%
0·030
Cl
0·05
0·026
21%
0·065
0·04
0·008
38%
0·077
0·07
0·057
12%
0·081
0·07
0·064
Total
98·98
95·15
1·31%
76·83
74·15
73·78
16%
0·27
1·24%
99·81
0·21
0·07
96·62
95·97
0·32%
74·62
19%
0·25
0·36%
97·51
76·42
76·14
0·13
0·12
98·62
98·11
0·31%
76·86
7%
0·15
0·33%
99·19
76·64
76·32
0·10
0·06
99·05
98·52
0·28%
76·94
24%
0·14
7%
0·078
0·32%
99·43
LA-ICP-MS analyses
n:
10
11
Rb
198
Sr
194·8
22·4
200·7
1·5%
Y
21·9
51·1
22·9
1·9%
Zr
48·9
225
52·7
2·25%
Ba
214·5
612
234·1
1·5%
195·1
498
222·1
18·4%
146·2
521
160·8
2·2%
143·5
507
153·6
1·6%
La
598·2
70·6
624·7
1·8%
358·1
69·3
583·3
5·1%
505·5
50·9
535·1
2·1%
496·7
49·0
520·3
2·9%
1·01%
10
211
199·4
15·4
4·24
48·8
45·9
213
3·26%
175
10
0·70%
177
0·85%
218·9
56·8%
172·0
11·5
175·8
2·8%
174·1
11·0
178·6
3·7%
27·1
3·5%
10·9
37·3
11·9
2·6%
10·2
35·7
11·5
2·8%
50·6
3·74%
35·3
154
38·4
2·50%
34·2
148
37·1
2·52%
Sm
69·0
10·44
73·0
3·5%
65·3
10·00
78·6
5·1%
48·8
7·54
52·4
7·4%
47·2
7·28
51·0
5·8%
Eu
9·84
0·94
10·89
5·3%
9·31
0·77
10·84
15·6%
6·48
0·48
8·26
6·3%
6·52
0·48
7·83
8·3%
Gd
0·85
8·72
1·06
7·1%
0·59
8·48
0·95
3·4%
0·44
6·54
0·53
6·1%
0·40
6·34
0·52
5·7%
Yb
7·67
5·30
9·99
3·8%
8·09
5·14
9·08
5·6%
5·91
3·76
7·08
5·6%
5·72
3·62
6·89
7·5%
Lu
4·91
0·75
5·61
4·0%
4·63
0·70
5·67
5·7%
3·51
0·51
4·21
3·9%
3·05
0·51
3·94
5·9%
Th
0·71
21·0
0·80
2·2%
0·62
19·3
0·77
3·4%
0·48
16·1
0·54
7·6%
0·47
15·2
0·57
2·6%
U
19·8
6·10
21·6
2·8%
18·2
5·47
20·3
1·3%
15·1
4·26
19·4
2·3%
14·6
4·28
15·8
2·1%
5·86
6·34
5·37
5·61
4·08
4·39
4·12
4·36
Values in bold indicate average concentrations. For each analysis, the per cent relative standard deviation is presented to
the right of average value (bold), and numbers below these indicate minimum and maximum concentrations measured.
n.a., not analyzed; n, number of analyses.
2108
GIRARD & STIX
35
1.2
(a)
30
Eu (ppm)
Sr (ppm)
25
20
15
10
0.8
0.6
0.4
0.2
5
0
0
0
35
200
400
600
Ba (ppm)
800
1000
150
1.2
(c)
30
170
190
210 230
Rb (ppm)
250
270
(d)
1
Eu (ppm)
25
Sr (ppm)
(b)
1
20
15
10
0.8
0.6
0.4
0.2
5
0
0
150
170
190
210 230
Rb (ppm)
250
270
0
5
10
15
20
Sr (ppm)
25
30
35
1000
(e)
Ba (ppm)
800
EBB (BS47)
EBB (BS89)
EBB (BS88-05)
EBB (BS87)
600
400
MBB
NBB
SBB
SL
Canyon Rhyolites
CF
TSC
DR (DR71-6)
DR (DR237B)
200
0
150
170
190
210 230
Rb (ppm)
250
270
Fig. 7. Glass compositions of Rb, Ba, Sr and Eu, which exhibit large variations in EBB samples BS89 and BS87, and TSC. Other glasses appear
homogeneous.
We observed strong variability in Rb (210^255 ppm), Sr
(5^20 ppm), Ba (300^700 ppm) and Eu (0·4^0·8 ppm) in
two EBB samples (BS89 and BS87), which define a trend
in which Rb correlates negatively with Sr, Ba and
Eu (Fig. 7), in continuity with more primitive and
homogeneous Ba, Sr and Rb concentrations for EBB samples BS47 and BS88-05. Rb is known to be sensitive to
alteration; several Central Plateau Member rhyolite samples substantially affected by devitrification show strong
Rb variability in the glass. In these samples, variability in
2109
VOLUME 50
other elements such as Sr, Ba and Eu is low and comparable with that for similar unaltered rhyolites with homogeneous Rb in the glass (Electronic Appendix 2). For altered
rocks, therefore, Rb is the sole heterogeneous element
observed. For the glassy EBB samples, we interpret the
observed variability of Rb, Sr, Ba and Eu as primary and
not related to alteration. TSC also shows similar variability
in Rb (200^220 ppm), Sr (4^27 ppm), Ba (360^580 ppm)
and Eu (0·6^1·0 ppm). TSC is overlain stratigraphically
by CF, and the ages of these units are indistinguishable
within error (Gansecki et al., 1996). CF exhibits compositional continuity with TSC for Rb, Sr and Ba, but its glass
is homogeneous and more primitive (Fig. 7). Other lavas
of the Biscuit Basin rhyolites (Table 1) have homogeneous
glass compositions that are typically more evolved than
those of EBB samples (e.g. 220^230 ppm Ba and
13^14 ppm Sr in NBB). In the Canyon rhyolites (Table 1),
DR stratigraphically overlies CF, and its age is similar to
that of CF and TSC (Gansecki et al., 1996). Nevertheless,
similar to our observations on mineralogy and whole-rock
geochemistry, no clear compositional link can be found
between DR and CF for the glass (Fig. 7), as concentrations in Ba, Sr and Eu on the one hand, and Rb on the
other hand, are both lower in DR than in CF. Data
obtained for two DR samples at two localities (Table 1;
Fig. 1) have similar compositions in all elements, ruling
out the possibility of post-eruptive alteration.
Furthermore, concentrations of nearly all major and trace
elements in glass are lower in DR than in CF or TSC, as
well as all other UBM rhyolites (Figs 5^8).
Although none of the samples are clearly heterogeneous
in Zr, this element shows a significant range of compositions for certain UBM glasses. Zr is highest in EBB at
255^295 ppm, whereas it is lower in MBB at 185^220 ppm,
in NBB at 165^175 ppm, and in SBB and SL at
140^150 ppm (Table 3; Fig. 8). In the Canyon rhyolites, Zr
varies from 215^235 ppm in CF and 195^220 ppm in TSC
to 145^160 ppm in DR. This range of variation for glasses
is comparable with that observed for the whole-rock analyses (Table 2, Fig. 4).
In contrast to the observed variability in Rb, Sr, Ba, Eu
and Zr, variation in other trace elements in the UBM rhyolite glasses is small and similar to that observed in the
whole-rocks (Fig. 8). For example, Y shows a variation of
45^55 ppm (except for DR, which has 35^40 ppm), La
65^80 ppm (47^52 ppm in DR), and Th 20^28 ppm
(14^17 ppm in DR). Variation within samples is limited,
and variation between samples is generally constant and
probably of analytical origin (Table 3; Electronic
Appendix 2).
A noticeable difference is observed between the Biscuit
Basin and the Canyon rhyolites for Th. At similar concentrations of Y, Th is systematically higher in the Biscuit
Basin glasses (Fig. 8). This results in distinct Y/Th ratios
NUMBER 11
NOVEMBER 2009
for the Canyon rhyolites (2·3^2·6) compared with the
Biscuit Basin rhyolites (2·0^2·2) (Fig. 8), suggesting that
the two groups of rhyolites are not cogenetic.
Plagioclase geochemistry
Representative analyses and anorthite contents for plagioclase are summarized in Table 4 and Fig. 9. For each
sample, the average, per cent relative standard deviation,
minimum and maximum anorthite and Ba contents are
presented (Table 4). Minimum contents of Ba are not
shown as they are near or below the instrumental detection
limit.
We observe no compositional difference within the
unsieved crystals between crystals (Fig. 2a) and the aggregates (Fig. 2b); the single crystals may thus simply be
either disaggregated or part of aggregates that are not
fully exposed in two-dimensional sections. Few analyses of
finely sieved crystals (Fig. 2d) were successful, although
these crystals appear compositionally similar to the coarsely sieved crystals. Compositions of unsieved and sieved
plagioclases are not identical, although they largely
overlap.
In East Biscuit Basin (EBB) sample BS89, unsieved plagioclase shows a large range of compositions from An23 to
An48 with an average of An32. Sieved plagioclase also
shows variability in composition, ranging from An24 to
An34 and averaging An28. EBB sample BS88-05 has
slightly more sodic plagioclase. In this sample, unsieved
crystals have An20^42 (average of An31), and sieved crystals
have An19^29 (average of An26), with no systematic compositional zoning in single crystals.
In the Middle Biscuit Basin flow (MBB), plagioclase is
slightly more sodic on average, although the calcic endmember values are higher than for EBB; unsieved crystals
exhibit a large range of An18^52 (average of An27), and
sieved crystals have An21^42 (average of An26) (Fig. 9). The
value of An52 is the most calcic plagioclase analysis documented for a Yellowstone rhyolite. In North Biscuit Basin
(NBB), unsieved plagioclase differs from EBB and MBB
by a lack of An contents 4An33, resulting in a lower average of An25. Sieved crystals have An contents similar to
those of MBB, with a range of An23^40 and an average of
An26. In SBB, An contents never exceed An28 and have an
average of An23, with unsieved crystals slightly more sodic
(An19^24) than sieved crystals (An22^28). In SL, An contents
never exceed An32 and have an average of An24, with no
clear compositional difference between the two plagioclase
groups (Fig. 9).
For the Canyon rhyolites, the most calcic plagioclase is
found in the Canyon flow (CF), where unsieved plagioclase shows An21^46 (average of An28) and sieved plagioclase An20^36 (average of An26). In the tuff of Sulphur
Creek (TSC), compositions are more sodic, with An20^32
(average of An26) for unsieved plagioclase and An20^29
(average of An24) for sieved plagioclase. The Dunraven
2110
GIRARD & STIX
1.6
350
(a)
300
1.2
Zr / Rb
Zr (ppm)
250
200
150
1
0.8
0.6
100
0.4
50
0.2
0
0
150
350
170
190
210 230
Rb (ppm)
250
270
150
7
(c)
300
250
5
200
4
150
190
2
50
1
40
50
60
30
40
Y (ppm)
270
50
60
Y (ppm)
4.5
350
(e)
300
(f)
4
250
3.5
Zr / La
Zr (ppm)
250
0
30
200
150
3
2.5
100
2
50
0
1.5
40
50
60
70
La (ppm)
80
90
40
50
60
70
La (ppm)
80
90
2.7
60
(g)
55
(h)
2.6
2.5
2.4
50
Y / Th
Y (ppm)
210 230
Rb (ppm)
3
100
0
170
(d)
6
Zr / Y
Zr (ppm)
(b)
1.4
45
40
2.3
2.2
2.1
2
35
1.9
30
1.8
10
15
20
Th (ppm)
25
30
10
15
20
Th (ppm)
25
Fig. 8. Glass compositions of selected trace elements and trace element ratios. Symbols are the same as for Fig. 7.
2111
30
VOLUME 50
NUMBER 11
NOVEMBER 2009
Table 4: Electron microprobe analyses of plagioclase
Unit:
Sample:
EBB
BS88-05
Mean
SiO2
TiO2
Al2O3
FeO
MgO
CaO
SrO
BaO
Na2O
K2O
Total
Cations
60·51
0·03
24·00
0·34
0·01
6·49
0·06
0·14
7·13
1·21
99·93
5·00
Plagioclase group
An
% RSD
Unsieved
Sieved, rims
Sieved, cores
n
Unsieved
Sieved, rims
Sieved, cores
Unsieved
Sieved, rims
Sieved, cores
Unit:
Sample:
31·1
(7·3)
25·9
(2·4)
26·3
(2·0)
1530
1870
1890
21
35
29
Max.
63·15
0·03
21·83
0·48
0·00
3·92
0·08
0·43
7·63
2·60
100·15
5·00
SiO2
TiO2
Al2O3
FeO
MgO
CaO
SrO
BaO
Na2O
K2O
Total
Cations
62·51
0·02
22·81
0·26
0·01
4·84
0·04
0·10
7·90
1·47
99·96
5·00
Plagioclase group
Unsieved
Sieved, rims
% RSD
Sieved, cores
% RSD
Ba (ppm) Unsieved
Sieved, rims
Sieved, cores
n
Unsieved
Sieved, rims
Sieved, cores
1
22·7
(1·4)
22·6
(0·7)
24·1
(1·6)
1140
1190
1080
23
18
22
Mean
57·95
0·05
25·96
0·31
0·01
8·69
0·07
0·05
6·20
0·67
99·95
5·00
41·9
20·2
28·9
18·8
29·0
b.d.l.
b.d.l.
b.d.l.
61·04
0·03
23·73
0·35
0·02
6·06
0·06
0·14
7·26
1·24
99·94
4·99
unsvd
20·4
3250
2640
2560
SBB
BS132
Mean
An
% RSD
Min.
svd (core)
% RSD
% RSD
Ba (ppm)
EBB
BS89
MBB
BM7-3
Min.
Max.
62·86
0·00
22·55
0·30
0·00
4·84
0·10
0·20
7·62
1·80
100·27
4·99
unsvd
31·5
(6·7)
27·9
(2·9)
27·6
(1·5)
1480
1570
1690
26
19
17
56·46
0·13
26·87
0·63
0·03
9·92
0·04
0·03
5·56
0·50
100·17
5·00
48·2
24·1
34·4
24·9
30·7
2370
2030
2360
SL
SL32-05
Min.
63·51
0·00
21·88
0·22
0·01
3·89
0·03
0·04
8·33
1·75
99·66
5·00
Max.
61·38 62·53
0·04
0·03
23·75 22·85
0·25
0·27
0·01
0·01
5·86
5·03
0·09
0·04
0·11
0·11
7·78
7·83
0·96
1·41
100·22 100·09
5·00
4·99
Svd (core)
18·5
21·6
21·7
b.d.l.
b.d.l.
b.d.l.
Mean
23·9
(1·2)
24·2
24·1
(2·5)
27·8
24·6
(0·4)
2460
1270
1940
1000
1720
1440
33
24
4
Min.
63·63
0·01
21·85
0·21
0·01
3·92
0·01
0·09
7·97
2·26
99·98
5·00
unsvd
27·4
(8·3)
27·7
(5·5)
25·3
(3·2)
1110
1120
1330
28
16
21
Min.
63·62
0·02
22·39
0·23
0·00
4·31
0·08
0·11
8·34
1·21
100·33
4·99
Max.
60·75
0·03
24·27
0·23
0·01
6·61
0·03
0·08
7·39
0·85
100·23
5·00
21·4
20·7
24·1
b.d.l.
b.d.l.
b.d.l.
25·9
Mean
54·80
0·04
28·18
0·28
0·01
11·11
0·01
0·00
5·22
0·40
100·06
5·01
Min.
62·06
0·03
23·10
0·29
0·01
5·28
0·03
0·10
7·48
1·46
99·85
4·99
unsvd
52·8
21·4
42·4
21·7
32·3
b.d.l.
b.d.l.
b.d.l.
Mean
2350
2320
1810
25·3
(3·1)
26·8
(4·5)
25·8
(2·3)
1090
1400
710
35
11
12
Min.
63·06
0·03
22·31
0·30
0·01
4·27
0·01
0·23
7·88
2·08
100·20
5·00
Svd (rim)
1
20·7
20·3
21·2
b.d.l.
b.d.l.
b.d.l.
Max.
56·65
0·07
26·51
0·33
0·02
9·52
0·04
0·13
5·73
0·55
99·55
4·99
Mean
61·99
0·03
22·92
0·31
0·01
5·02
0·06
0·19
7·61
1·61
99·75
5·00
Svd (rim)
46·3
25·9
(4·2)
32·2
23·3
(2·7)
35·5
24·0
(1·6)
3080
2070
2930
2310
3070
1990
19
28
15
Max.
62·98
0·02
22·09
0·30
0·02
4·20
0·01
0·18
7·47
1·98
99·27
4·97
58·75
0·02
25·89
0·25
0·01
8·16
0·03
0·06
6·36
0·71
100·23
4·99
unsvd
svd (rim)
20·9
33·1
24·0
39·8
22·9
29·1
b.d.l.
b.d.l.
b.d.l.
TSC
TSC217-2
61·94
0·03
23·07
0·31
0·01
5·50
0·06
0·17
7·46
1·43
99·97
4·99
27·5
(5·9)
31·5
25·3
(3·4)
25·0
26·0
(4·1)
1900
1880
1450
1960
1680
1870
37
21
18
Max.
18·6
CF
CF133B
Svd (rim) Svd (rim)
24·3
61·71
0·03
23·39
0·30
0·01
5·58
0·04
0·11
7·59
1·46
100·21
5·00
unsvd
23·3
b.d.l.
b.d.l.
b.d.l.
Mean
NBB
BN154A
2460
2330
1490
DR
DR237B
Min.
Max.
Mean
Min.
63·43 60·53 62·95 64·00
0·01
0·01
0·02 0·01
22·08 24·26 22·50 21·90
0·28
0·30
0·33 0·29
0·01
0·01
0·01 0·00
4·10
6·60
4·59 3·78
0·08
0·07
0·03 0·04
0·27
0·14
0·14 0·21
7·82
7·08
8·00 8·32
2·15
1·06
1·54 1·82
100·23 100·06 100·09 100·35
4·99
4·99
5·00 5·00
1
unsvd
19·7
31·9
19·7
29·3
22·0
(1·5)
—
22·1
27·5
—
b.d.l. 3460
b.d.l. 3540
b.d.l. 2580
1600
Max.
62·20
0·00
23·07
0·30
0·01
5·11
0·00
0·13
7·79
1·31
99·93
4·99
unsvd
18·0
24·6
—
—
—
—
b.d.l. 2580
32
Minimum and maximum analyses selected are the analyses with the minimum and maximum An content for the whole
sample, regardless of their plagioclase group. ‘Plagioclase group’ field indicates which plagioclase group these analyses
belong to (unsvd, unsieved, svd, sieved). For An contents, per cent relative standard deviations are indicated in parentheses. These are an indication of the range of composition. b.d.l., below detection limit; n, number of analyses.
2112
GIRARD & STIX
70
(a)
60
Canyon Rhyolites
70
EBB
60
50
40
40
CF
n
n
50
(b)
30
30
20
20
10
10
0
16
70
60
20
24
28
32
(c)
36
An
40
44
48
52
0
56
16
70
MBB
60
50
40
40
24
28
32
(d)
36
An
40
44
48
52
56
40
44
48
52
56
40
44
48
52
56
TSC
n
n
50
20
30
30
20
20
10
10
0
16
70
60
20
24
28
32
(e)
36
An
40
44
48
52
0
56
16
70
NBB
60
50
40
40
24
28
32
(f)
36
An
DR
n
n
50
20
30
30
20
20
10
10
0
16
70
60
20
24
28
(g)
32
36
An
40
44
48
52
56
0
16
20
24
28
32
36
An
SBB + SL
50
Sieved
30
Unsieved
n
40
20
10
0
16
20
24
28
32
36
An
40
44
48
52
56
Fig. 9. Plagioclase compositions in various rhyolite units, shown for unsieved vs sieved crystals.
2113
VOLUME 50
NUMBER 11
NOVEMBER 2009
Fig. 10. Photomicrograph of a zoned unsieved plagioclase from unit EBB (sample BS88-05) with high-Ba and low-An rims. View under crossed
polars, through a 100 mm thick section.
Road flow (DR) again stands out from the other UBM
rhyolites by having the most sodic plagioclase with
An18^25 (average of An22) (Fig. 9).
Despite the large range of An contents found in crystals
of many UBM rhyolites, there is no clear systematic coreto-rim zoning for unsieved crystals, nor between cores
and fresh rims of sieved crystals. This lack of systematic
zoning may reflect complex small-scale oscillatory zoning
that was not revealed at the resolution of our core-to-rim
profiles. For unsieved crystals of EBB sample BS88-05,
however, all low-An (5An24) values are found on rims,
whereas all high-An values (4An24) are found in cores
(Fig. 10). This could not be verified for other samples
because of the limited number of high-An and low-An
analyses obtained, and the limited number of analyses
obtained on each grain.
As predicted thermodynamically (Seck, 1971; Ribbe,
1975), the K2O content in plagioclase decreases with
increasing An content. At a given An content, however,
the K2O content increases with temperature (Seck, 1971;
Ribbe, 1975). No appreciable differences in K2O were
found between unsieved and sieved plagioclase for any of
the UBM rhyolites. However, changes in K2O for a given
An content (An23) are observed between the units. K2O is
highest in EBB with 1·7^1·8 wt% at An23. K2O is slightly
lower in MBB and NBB at 1·6^1·8 wt%, and lowest in
SBB and SL at 1·4^1·6 wt% (Fig. 9). For the Canyon
rhyolites, K2O at An23 is 1·6^1·7 wt% in CF and TSC and
lower at 1· 4^1·5 wt% in DR (Fig. 11).
The distribution of Ba in plagioclase also decreases with
increasing An content (Blundy & Wood, 1991); only at low
An content was Ba sufficiently high to be measured with
reliability. As a result, minimum Ba values measured are
not listed in Table 4. High-Ba analyses (41800 ppm) were
typically found on unsieved plagioclase rims, with highBa rims particularly well developed for unsieved crystals
of EBB sample BS88-05 (Fig. 10). In other UBM rhyolites
such as MBB, NBB, CF and TSC, a few high-Ba analyses
were obtained. In these samples, however, the low-An
analyses were not systematically Ba-rich, possibly reflecting crystallization from a melt with lower Ba. In SBB and
SL, Ba is significantly lower (average of 1000^1400 ppm,
maximum values typically 52000 ppm); plagioclase in
these samples is more sodic, and increased contents of Ba
in plagioclase are thus to be expected. It is thus possible
that the parental melt was again lower in Ba.
DISCUSSION
A single magmatic series or independent
magma batches?
Observed trends in whole-rock and plagioclase geochemistry appear to reveal a cogenetic origin for the Upper
Basin Member (UBM) rhyolites. For example, CaO,
2114
GIRARD & STIX
K2O (wt %)
3
An23
EBB
MBB
SBB & SL
DR
2
1
0
15
20
25
30
35
40
45
50
55
An content
Fig. 11. Variation of An content vs K2O for representative Upper Basin Member rhyolites.
MgO, FeO, Sr and Eu generally decrease with SiO2
(Fig. 4), although MgO and TiO2 suggest more complex
or several fractionation trends. In parallel, plagioclase
becomes progressively more sodic, indicating progressively
cooler crystallization environments (Figs 9 and 11). These
observations are consistent with the observed changes in
mineralogy, which show a progressive increase in sanidine
and quartz abundances at the expense of plagioclase
(Fig. 3).
Many high-silica rhyolite sequences such as largevolume ignimbrites have been interpreted as originating
from a single zoned magma chamber. In these systems,
progressive decreases in crystallinity, temperature, Sr, Ba,
LREE, etc., are typically observed towards the base of the
ignimbrite, together with increases in many incompatible
trace elements such as Rb, Th, U, Y, Nb, etc., by a factor
of two or more (e.g. Hildreth, 1981; Mahood, 1981;
Hildreth et al., 1984; Streck & Grunder, 1997; Anderson
et al., 2000; Hildreth & Wilson, 2007).
In contrast, no clear trends are observed in the UBM
rhyolites as a function of Ba and Sr for any of the trace elements mentioned above. The ranges of composition in the
whole-rock and the glass are small (Tables 2 and 3), typically not greater than 30% of the lowest measured concentrations. Therefore, we propose that the UBM rhyolites
may instead originate from several independent magma
batches that initially originated from a common source
and later evolved by fractional crystallization and magma
mixing, possibly with a silicic replenishing magma. We
propose that at least six magma batches were involved in
the formation of the Upper Basin Member rhyolites. We
now examine evidence for this hypothesis, based on trace
element distributions that rule out an origin from a single
petrogenetic suite in a single magma chamber.
Southwestern Biscuit Basin rhyolites vs Northeastern
Canyon rhyolites
These two eruptive foci are essentially coeval, with the
Middle Biscuit Basin flow (MBB) in the SW dated at
516 7 ka and the Canyon flow (CF), tuff of Sulphur
Creek (TSC) and Dunraven Road flow (DR) in the NE
dated at 484 15, 473 9 and 486 42 ka, respectively
(Gansecki et al., 1996) (Table 1; Fig. 1). Nevertheless, the
two eruptive foci are 50 km apart, and it is therefore
questionable whether they originate from the same
magma chamber. Rhyolites in each geographical group
encompass a wide range of mineralogy and major element
compositions but have generally similar trace element geochemistry. Some exceptions in this trace element distribution are clear, however, indicating that the two groups
cannot be cogenetic.
Regardless of composition and degree of evolution, the
Y/Th ratio in glass clearly discriminates the Canyon rhyolites from the Biscuit Basin rhyolites (Fig. 8), with Y/Th
ratios of 2·3^2·6 observed for the Canyon rhyolites compared with 2·0^2·2 for the Biscuit Basin rhyolites. Values
for Th in whole-rocks are similar for both groups
(Table 2). Therefore, our glass data suggest that different
types and/or abundances of accessory mineral phases control the distribution of Th, rather than a different source.
Concentrations of Th in the Biscuit Basin rhyolite glasses
approach or slightly exceed those of the whole-rocks,
whereas the Canyon rhyolite glasses are clearly depleted
in Th compared with the whole-rocks (Tables 2 and 3).
Trace element analyses of zircon from Bindeman & Valley
(2001) and Bindeman et al. (2008) do not clearly show
higher Th in the Canyon rhyolite zircons. Also, the glass/
whole-rock ratios of Zr are similar in the two rhyolite
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VOLUME 50
groups, ruling out differences in the amount of zircon crystallizing. The Canyon rhyolites may thus contain another
Th-bearing accessory mineral phase. Regardless of the
mineral responsible for the observed Th depletion in the
Canyon rhyolites glasses, our data show different behaviour of Th at each focus of the early post-caldera magmatism, requiring that the magmas that generated the
Canyon rhyolites evolved separately from those that generated the Biscuit Basin rhyolites. We now examine details
of magma evolution within the Canyon rhyolites and the
Biscuit Basin rhyolites.
The unusual composition of the Dunraven Road lava flow
This lava flow exhibits five characteristics that make it
highly unusual among the Canyon rhyolites (and other
UBM rhyolites). (1) It has low crystallinity (52%), contrasting with the 10^20% crystals generally observed in
other UBM rhyolites (Table 2, Fig. 3). (2) Plagioclase is
slightly more sodic than in other UBM rhyolites and has
lower K2O at similar An contents, suggesting that the DR
magma evolved under lower temperatures (Table 4,
Fig. 9). (3) It has the highest SiO2 among all UBM rhyolites, contrasting with its relatively primitive mineralogy
in which plagioclase is by far the dominant species,
whereas quartz and sanidine are not found in appreciable
amounts. (4) Geochemically, it is depleted in nearly all
trace elements in both whole-rock and glass relative to
other UBM rhyolites (Figs 4^8). Whole-rock concentrations for most analyzed trace elements are 20% lower
than in other UBM rhyolites; for example, its REE pattern
is distinct from those of other UBM rhyolites, which are
generally indistinguishable from each other (Fig. 6).
(5) Compared with other elements, DR has a more pronounced negative Eu anomaly (Fig. 6) and more pronounced depletion in Sr and Ba relative to other UBM
rhyolites (Figs 4, 7 and 8).
These differences demonstrate that DR is compositionally distinct from the other northeastern Canyon rhyolites,
namely the Canyon flow (CF) and tuff of Sulphur Creek
(TSC). These geochemical differences are clearly visible
in Figures 7 and 8, and we propose that DR represents a
magma batch distinct from CF and TSC. A higher degree
of partial melting for DR may explain the depletion in
trace elements such as Rb, Y, Th, U, REE, etc., and it is
likely that feldspar fractionation subsequently occurred, as
DR is also depleted in Sr, Ba and Eu. Another possibility
is that DR may result from bulk melting of a source
depleted in large ion lithophile elements (LILE), high
field strength elements (HFSE) and REE compared with
the source(s) of the other UBM rhyolites.
NUMBER 11
NOVEMBER 2009
EBB to 15^20 ppm in South Biscuit Basin (SBB) and
Scaup Lake (SL), and Ba decreasing from 740^780 ppm
in EBB to 360^380 ppm in SBB and SL. Europium in
glass also is lower in SBB and SL (Fig. 7). In the wholerock, Sr and Eu also are lower in SBB and SL (Figs 5 and
6). These depletions could be explained by feldspar fractionation in a closed system. However, in such a scenario,
Rb should increase, similar to what is observed in many
zoned ignimbrites (e.g. Hildreth, 1981; Streck & Grunder,
1997; Anderson et al., 2000). No clear increase of Rb is
observed from EBB to SBB and SL, neither in the wholerock nor in the glass. Furthermore, many zoned rhyolitic
units exhibit a pronounced decrease in LREE together
with Sr and Ba (e.g. Hildreth, 1981; Streck & Grunder,
1997). Here, by contrast, there is a small increase in
LREE in the whole-rock (Fig. 6), resulting in a Ce/Yb
ratio that is higher in SBB and SL than in any other
UBM rhyolite (Fig. 5). Therefore, we propose that SBB
and SL are not cogenetic with the primitive EBB rhyolites.
The South Biscuit Basin and Scaup Lake lava flows also
are much younger than the other dated Upper Basin
Member rhyolites (Table 1). SBB has been dated at
255 11 ka (Bindeman et al., 2008) whereas SBB has been
dated at 257 13 ka (Christiansen et al., 2007). The older
UBM rhyolites are low-d18O rhyolites, with quartz d18O
values of 0^1ø SMOW, contrasting with the values of
4^5ø observed in SBB and SL (Hildreth et al., 1984;
Bindeman & Valley, 2001). EBB also appears to be a lowd18O rhyolite (Sturchio et al., 1986), although its age has
not been determined. These changes in the isotopic signature imply that SBB and SL cannot have been derived
from the low-d18O UBM rhyolites solely by closed-system
magma chamber processes. Based on d18O zoning of zircons, Bindeman & Valley (2001) proposed that the SBB
and SL zircon cores were partially inherited from these
early post-caldera, low-d18O magmas. We propose that
SBB and SL may result from the reactivation of the magmatic system after a prolonged repose period, rather than
from magmas that evolved from EBB. Mixing of such
low-d18O magmas with magma significantly less depleted
in 18O may have led to the formation of a new eruptible
magma body at 250 ka. Replenishing silicic magmas
constitute a possible mixing component. SBB and SL may
in fact represent precursors for the younger voluminous
Central Plateau Member rhyolites emplaced between 174
and 70 ka (Christiansen, 2001; Christiansen et al., 2007).
In this scenario, replenishing magmas may have played
an important role in the generation of of SBB and SL. The
relationship between SBB and SL and the Central Plateau
Rhyolites is dealt with in a separate publication (Girard
& Stix, in preparation).
The South Biscuit Basin and Scaup Lake lava flows
These two lava flows have lower Sr and Ba in glass than
the primitive East Biscuit Basin (EBB) samples BS88-05
and BS47 (Fig. 7), with Sr decreasing from 25^30 ppm in
The North Biscuit Basin lava flow
A third Biscuit Basin rhyolite, the North Biscuit Basin lava
flow (NBB, Table 1) has intermediate d18O values of 3^5ø
2116
GIRARD & STIX
(Bindeman & Valley, 2001), and therefore cannot be
derived from the East Biscuit Basin (EBB) or Middle
Biscuit Basin (MBB) magmas by closed-system magma
chamber processes. Its d18O signature is similar to that of
South Biscuit Basin (SBB) and Scaup Lake (SL) flows
(Bindeman & Valley, 2001), but the following characteristics suggest that NBB is not cogenetic with SBB and SL.
When compared with these two lava flows, NBB has a
lower Ce/Yb ratio similar to those of EBB and MBB,
higher Zr/Y and Zr/La ratios (Fig. 8d^f), and lacks the
large pyroxene phenocrysts and the low FeO glass that are
characteristic of both SBB and SL. NBB also contains mineral aggregates similar to EBB and MBB. No age has been
determined for NBB, and its relationship to other Upper
Basin Member rhyolites is unclear.
The Middle Biscuit Basin lava flow
The Middle Biscuit Basin lava flow (MBB, Table 1) is a
low-d18O rhyolite with plagioclase as the dominant mineral, which occurs either as sieved crystals or in aggregates
similar to those in the East Biscuit Basin (EBB) flow.
Among the Biscuit Basin rhyolites, this unit appears to be
the most primitive lava apart from EBB. It has lower Sr
and Ba than EBB, which could suggest that it is derived
from EBB by feldspar fractional crystallization. However,
the Rb content is not significantly higher than that of the
EBB, which is opposite to that expected from fractional
crystallization or melting of feldspar (see below). It also
has a different signature in terms of Zr (Fig. 8), hence may
represent a magma batch separate from the EBB.
Eu, lower Rb), although also maintaining compositional
continuity with samples BS89 and BS87 (Fig. 7). Similarly,
the homogeneous Canyon flow (CF) sample CF133B continues the trend observed for TSC and is more primitive
(Fig. 7). Based on modeling results, we compare the likelihood of fractional crystallization vs magma mixing and
melting to generate these trends.
Assuming that fractional crystallization occurs in a
closed-system magma chamber, the change in trace element concentration is described by
CL
¼ F ðD1Þ
C0
ð1Þ
where CL and C0 are the trace element concentrations in
the final and the initial liquids, respectively (here, our
best estimates of these parameters are the evolved and
primitive end-member concentrations, respectively), F is
the volume fraction of residual melt, and D is the bulk
crystal/melt distribution coefficient of each trace element.
Equation (1) can be reformulated as
1=ðD1Þ
CL
:
ð2Þ
F¼
C0
For each trace element, the bulk distribution coefficient D
corresponds to
pl
qz
ð3Þ
pl
qz
ð4Þ
þ Xqz KDBa
DBa ¼ Xpl KDBa þ Xsan KDsan
Ba
þ Xqz KDSr
DSr ¼ Xpl KDSr þ Xsan KDsan
Sr
pl
qz
DRb ¼ Xpl KDRb þ Xsan KDsan
þ Xqz KDRb
Rb
Summary
In summary, we propose the existence of six independent
magma batches that formed the UBM rhyolites. Four
magma batches, associated with five lava flows, were present in the SW near the Mallard Lake Resurgent Dome:
(1) East Biscuit Basin (EBB), (2) Middle Biscuit Basin
(MBB), (3) North Biscuit Basin (NBB) and (4) South
Biscuit Basin (SBB) and Scaup Lake (SL), which cannot
be related to one another based on our geochemical data.
Two batches, associated with three rhyolite units, were
present in the NE near the Sour Creek Resurgent Dome:
(5) Dunraven Road (DR) and (6) tuff of Sulphur Creek
(TSC) and Canyon flow (CF). We now discuss the petrogenesis of the EBB samples and the TSC^CF succession.
Compositional variability in East Biscuit
Basin and tuff of Sulphur Creek glass
Large variations are observed for Ba, Sr, Eu and Rb in
glass in two East Biscuit Basin (EBB) samples, BS89 and
BS87, as well as a tuff of Sulphur Creek (TSC) sample
TSC217-2 (Table 3; Fig. 7). Essentially, Ba, Sr and Eu
decrease whereas Rb increases. The other EBB samples
BS88-05 and BS47 are homogeneous for these elements
with slightly more primitive compositions (higher Sr, Ba,
ð5Þ
where X is the fraction of each mineral that crystallizes
such that Xpl þ Xsan þ Xqz ¼1. The distribution coefficient
KD represents the partitioning of an element between
liquid and a mineral phase, with the subscripts pl, san and
qz referring to plagioclase, sanidine and quartz, respectively. Methods and reasoning for determining the different KD values are detailed in Electronic Appendix 3, with
values used for modeling in Table 5.
The trend from the primitive EBB samples to evolved
EBB samples is satisfactorily explained for Sr, Ba and Rb
by 25% fractional crystallization, of which 20% is plagioclase, 30% sanidine and 50% quartz, whereas the
CF^TSC trend is satisfactorily explained for Sr and Ba by
15% fractional crystallization, of which 40% is plagioclase, 30% sanidine and 30% quartz (Table 5). For EBB,
the lack of quartz and sanidine in the mineral assemblage
is not consistent with the model predictions, although the
removal of newly formed crystals by settling remains possible. However, if quartz and sanidine were removed, the
presence of large sieved plagioclase crystals and denser
pyroxene^plagioclase aggregates in the EBB mineral
assemblage is problematic. Similarly, plagioclase is dominant over quartz and sanidine in both TSC and CF.
2117
VOLUME 50
NUMBER 11
NOVEMBER 2009
Table 5: Results of modeling of fractional crystallization and melting in EBB and TSC samples
East Biscuit Basin
Tuff of Sulphur
Creek–Canyon flow
Cevolved (ppm) (CL)
Ba
300
Sr
5
300
Rb
255
220
Ba
780
620
Sr
30
23
Rb
200
195
4·5
Cprimitive (ppm)
Partition coefficients (KD)
Ba
Sr
Rb
2
2
san
13
13
pl
18
18
san
12
12
pl
pl
0·25
0·25
san
0·5
0·5
FC
M
FC
M
Ba
780
880
620
736
Sr
30
81
23
71
Rb
200
185
195
172
C0 (ppm)
F
0·75
(1 – F)
0·25
Xpl
0·19
0·9
0·4
0·98
Xsan
0·31
0·1
0·3
0·02
Xqz
0·5
0
0·3
0
DBa
4·4
3·1
4·7
DSr
7·1
DRb
Comment
0·2
0·075
0·85
—
0·15
17
10·8
0·28
0·25
0·13
—
2·2
18
0·25
No quartz or
Modeling problematic
Less quartz or
sanidine observed
for Rb (bulk D
sanidine observed
for Rb (bulk D of
in final
of 0·7 required)
in final assemblage
0·75 required)
assemblage
Modeling problematic
plagioclase dominant
Cevolved and Cprimitive correspond to the lowest and highest concentrations, respectively, of Ba and Sr in glass
measured in samples from each series. C0 represents the composition of the initial liquid, in the case of
fractional crystallization (FC), and that of the initial solid, in the case of melting (M). For fractional crystallization, C0 is assumed to be the composition of the primitive glass, Cprimitive. For melting, we obtain C0
assuming that Cprimitive ¼ CL (the composition of the liquid) at a degree of melting that corresponds to the
amount of glass present in the whole-rock (see text for discussion). F is the fraction of melt, (1–F) the
amount of crystals formed, X the relative abundance of each mineral that is crystallizing or melting (pl,
plagioclase; san, sanidine; qz, quartz) and D is the bulk distribution coefficient of each element based on
this mineral assemblage [see equations (3)–(5)]. More detailed comments on the results of each model and
our preferred model are given in the text. Details on the selection of the partition coefficients are given in
Electronic Appendix 3 (KD for all studied elements in quartz is assumed to be zero).
We now examine mixing of two magmas as an alternative explanation for the observed compositional heterogeneity. In such a scenario, one magma has a composition
in Sr, Ba and Rb that corresponds to that observed in the
evolved glass, whereas the second magma, perhaps a
replenishing magma, has a composition in Sr, Ba and Rb
that corresponds to that observed in the primitive glass.
This replenishing magma may have had a signature in elements such as REE and HFSE similar to that of the fractionating magma, as no bimodal compositions or large
2118
GIRARD & STIX
compositional scatters are observed within the EBB
glasses. Because TSC is a pyroclastic deposit, its explosive
eruption may have initiated mixing of two magmas of
distinct compositions that were stored in one stratified
reservoir, as recently reproduced experimentally by
Kennedy et al. (2008). However, to explain the origin of
such a gradient in the magma chamber, fractional crystallization would be required.
A third possibility to explain these compositional variations is melting. Equilibrium melting is described by
CL
1
¼
C0 F þ D FD
ð6Þ
where CL and C0 are the trace element concentrations in
the liquid and the protolith, respectively, F is the volume
fraction of melt formed, and D is the bulk crystal/melt distribution coefficient of each trace element. Therefore, F is
described by
C0
D
CL
:
ð7Þ
F¼
1D
An obstacle to such modeling is estimating C0, the composition of the protolith. Assuming that the primitive EBB
samples and CF represent the samples formed by the highest degree of melting of each of the EBB and TSC^CF
suites, it is possible to estimate C0 given that F, the volume
fraction of melt (now the volume fraction of glass observed
in the rock) and CL (now the glass composition) are
known. The mineralogy of the samples is a good indicator
of the nature of the mineral assemblage that was melting
when the lavas were erupted. EBB samples BS88-05 and
BS47 essentially comprise plagioclase as minerals and
have 90% glass (hence F ¼ 0·9). Thus, we can calculate
C0 to be 880 ppm for Ba, 80 ppm for Sr and 185 ppm
for Rb (Table 5). Similarly, for CF, which comprises
85% glass and small amounts of sanidine and quartz in
addition to plagioclase, we estimate C0 to be 735 ppm
for Ba, 72 ppm for Sr and 172 ppm for Rb (Table 5).
The evolved end-member composition of EBB is modeled
satisfactorily for Ba and Sr by 7·5% melting of an assemblage of C0 composition comprising 90% plagioclase
and 10% sanidine, whereas the evolved end-member
composition of TSC is modeled by 13% melting of plagioclase in a source of C0 in composition (Table 5). For both
trends, however, a problem persists with Rb, for which
the concentration in the evolved end-member glass
(255 ppm for EBB and 220 ppm for TSC) is much lower
than predicted using KDRbpl ¼ 0·25 (560 ppm for EBB and
490 ppm for TSC). For the concentrations in Rb observed,
bulk distribution coefficients of 0·70 and 0·75 are
required in the EBB and TSC magmas. These values are
much higher than those commonly accepted for plagioclase in silicic magmas (e.g. Ren et al., 2003). On the other
hand, the mineral assemblage and textures on crystals in
EBB and TSC are in better agreement with a melting
model. The models predict that plagioclase is the sole or
dominant mineral phase that is melting. This agrees well
with the lack of sanidine and quartz in EBB, and the abundance of plagioclase in both EBB and TSC^CF.
Furthermore, plagioclase in both EBB and TSC^CF is
sieved; therefore, these heterogeneous glass compositions
may reflect the continuing disssolution of these crystals.
Because heterogeneity is developed at a sub-millimetre
scale, it is possible that these elements had not yet diffused
into the glass when the magma was erupted. This probably
implies that the melting event occurred soon before eruption. We now discuss in more detail the significance of the
sieve textures and the coexistence of sieved crystals and
aggregates for the petrogenesis of the UBM rhyolites.
Disequilibrium textures
Sieve textures in plagioclase
Perhaps the most spectacular feature of the UBM rhyolites
is the abundance of sieved plagioclase crystals. In the
third Yellowstone caldera cycle, these features are essentially restricted to the UBM rhyolites. Sieve textures in plagioclase have been obtained by experimental melting
(Johannes, 1989), and melting is the most commonly
accepted hypothesis to explain their origin (e.g. Keller,
1969; Genc alioglu Kuscu & Floyd, 2001; Bachmann et al.,
2002; Barbey et al., 2005). Here, melting is the process that
best explains the compositional variability of the glass in
the East Biscuit Basin and tuff of Sulphur Creek samples.
Therefore, a fundamental question concerns the cause of
the melting process.
The last pulse of volcanism at Yellowstone occurred at
70 ka; the current state of the magma chamber is interpreted to be a semi-solid non-eruptible crystal mush
(Miller & Smith, 1999; Husen et al., 2004). It is therefore
reasonable to infer that the Yellowstone magma reservoir
was also a crystal mush during the 100^150 kyr interval
after caldera collapse at 640 ka, during which most UBM
rhyolites were emplaced. Experiments on magma replenishment in a crystal mush have shown that a buoyant
replenishing silicic magma is capable of entraining up to
12% of silicic mush crystals (Girard & Stix, 2009); rhyolites may therefore comprise a mix of remobilized crystals
and new replenishing melt. In this scenario, the mush crystals are likely to undergo a sudden temperature increase,
leading to melting following their transfer into the replenishing magma. During ascent, adiabatic decompression
also will occur. Such decompression is also known to
favour the development of sieve textures in plagioclase
(Nelson & Montana, 1992), as well as to enhance the melting process.
A majority of the UBM rhyolites have a strongly
d18O-depleted signature. It is well established that
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VOLUME 50
magmas unaffected by hydrothermal interaction at
Yellowstone (e.g. the Huckleberry Ridge Tuff, the ignimbrite associated with the first caldera, as well as the extracaldera rhyolites) have normal magmatic d18O values of
7^8ø (Hildreth et al., 1984; Bindeman & Valley, 2001).
Therefore, the protolith of the UBM rhyolites probably
consists mainly of previously hydrothermally altered material. Bindeman & Valley (2001) and Bindeman et al.
(2008) have suggested bulk melting of such material. In
such a scenario, the sieved plagioclase may represent remnants of the major mineral phases of the hydrothermally
altered crystal mush or pluton.
Mineral aggregates
Another unusual aspect of the UBM rhyolites is the ubiquitous presence of small aggregates consisting of plagioclase,
clinopyroxene, orthopyroxene and Fe^Ti oxides. An
important question is whether these relatively mafic aggregates (60% SiO2, calculated from mass balance of crystal
compositions and abundances) were in equilibrium with,
and crystallized from, a high-SiO2 rhyolitic melt.
Because of their bulk composition, the aggregates could
represent either microdiorite xenoliths or crystals from an
andesitic to dacitic replenishing melt. Compositions of
both plagioclase and clinopyroxene, however, are typical
of those observed in rhyodacites and rhyolites (e.g.
Chesner, 1998; Barbey et al., 2005; Smith et al., 2005)
rather than those observed in andesites or dacites (e.g.
Murphy et al., 2000; Izbekov et al., 2002; Cooper & Reid,
2003; Zellmer et al., 2003). Additionally, the compositions
of plagioclase in the aggregates generally overlap those of
the sieved plagioclase, differing from these only by the
occasional presence of zones with An40. An origin as
cumulates from a silicic magma chamber is another possibility. However, this hypothesis is not consistent with the
general freshness of the crystals (which lack dissolution
textures), their fine-grained nature, and the lack of sanidine and quartz.
Similar aggregates are found in the Were Ilu ignimbrite
of Ethiopia (Barbey et al., 2005), a high-silica rhyolite that
also has abundant sieved plagioclase and lacks sanidine
and quartz, similar to the East Biscuit Basin (EBB) lavas.
These aggregates have been interpreted as the result of
assimilation^fractional crystallization of a rhyodacitic
magma. For the Were Ilu system, single plagioclase crystals
in the aggregates are normally zoned, with some cores
attaining or exceeding An50. Clinopyroxene in the Were
Ilu aggregates has a composition of Wo36En32Fs32, similar to that of clinopyroxene in the UBM rhyolites
(Hildreth et al., 1984; this study). Interstitial glass in these
aggregates and melt inclusions in plagioclase indicate crystallization from a melt with 70^74% SiO2 (Barbey et al.,
2005). It is thus possible that the UBM aggregates also
NUMBER 11
NOVEMBER 2009
crystallized from a rhyodacitic or a low-SiO2 rhyolitic
magma, potentially a primitive replenishing magma.
However, zoning in plagioclase is less common and less
pronounced than for the Were Ilu ignimbrite, and despite
the low-silica bulk composition of the EBB, there is no
direct evidence in EBB and other UBM rhyolites of a
replenishing 70^74% SiO2 magma. Yet glass in EBB has a
lower silica content (576·5% SiO2), which may be evidence of a minor contribution of such a replenishing
magma to a normal Yellowstone high-silica rhyolitic
magma. If 15% by volume of a replenishing magma
with 74% SiO2 is mixed with 85% of a crystal mush
with 77% SiO2, the product of mixing in these proportions is 76·5% SiO2, similar to the composition of the
EBB glass.
As it interacts with this resident rhyolitic crystal mush or
pluton, the replenishing magma probably exchanges heat
and cools, so that fractional crystallization may begin.
The first minerals formed are Fe^Ti oxides, pyroxenes
and plagioclase (Scaillet et al., 1995; Dall’Agnol et al., 1999).
We propose that the aggregates are formed at this stage.
As the replenishing magma cools, the mush temperature
increases and melting is enhanced, leading to the formation of sieve textures in the resident plagioclase and heterogeneous compositions in the glass. In such a scenario, the
plagioclase in the aggregates would have higher d18O
values than the sieved plagioclase (which would have a
true low-d18O signature). However, the expected higher
d18O signature of the aggregate crystals, which would be
clear evidence for replenishment, may be erased by fast diffusion of oxygen in feldspar (Bindeman & Valley, 2001),
making the validation of this hypothesis by the use of
oxygen isotopes difficult.
In other UBM rhyolites of more evolved composition,
where the glass ranges from 76·7 to 77·7% SiO2, the contribution of the replenishing magma may be lower, resulting
in a more subtle compositional change. The observed
decrease of the aggregate abundances with the increase in
quartz and sanidine throughout the UBM rhyolites is of
fundamental importance. In these rhyolites, the quartz
crystals generally appear to be in dissolution, suggesting
that they share a common origin with the sieved plagioclase, rather than with the aggregates. We hypothesize
that smaller-volume replenishments of this hot, more primitive magma occurred in the magma reservoirs at the
origin of the Middle Biscuit Basin, North Biscuit
Basin, Canyon flow and tuff of Sulphur Creek units,
allowing for some quartz and sanidine from the mush to
be preserved and fewer aggregates to crystallize. In the
magma reservoirs forming the South Biscuit Basin, Scaup
Lake and Dunraven Road lava flows, still less replenishment occurred, and no aggregates were formed at this
stage.
2120
GIRARD & STIX
A MODEL FOR T H E OR IGI N OF
THE U PPER BASI N MEMBER
R H YO L I T E S
Based on our observations and data, as well as available
age and isotope geochemistry information, we propose
that the Upper Basin Member magmatism evolved as follows. The third Yellowstone caldera collapse occurred at
639 2 ka (Lanphere et al., 2002), evacuating 1000 km3
of high-silica rhyolite magma (the Lava Creek Tuff).
Much of this volume was deposited within the caldera as
ignimbrites with thicknesses of several hundred metres
(Christiansen, 2001). This material was probably emplaced
on top of intracaldera ignimbrite from the 2·0 Ma
Huckleberry Ridge Tuff (Christiansen, 2001), increasing
the thickness and volume of the intracaldera rhyolites
(Fig. 12a).
Deep magma injections, probably of mafic composition
and related to the Yellowstone hotspot, stalled beneath the
caldera, resulting in uplift of two resurgent domes within
the caldera, the Mallard Lake and Sour Creek Resurgent
Domes (Fig. 12b). This uplift caused fracturing throughout
the caldera, leading to the development of a large hydrothermal system. This hydrothermal system was most developed beneath the resurgent domes, owing to maximum
uplift, fracturing, and the presence of magmatic heat
sources at depth. As a result, the intracaldera ignimbrite
acquired a low-d18O signature, which was particularly pronounced beneath the domes. Hydrothermal alteration and
d18O depletion may also have affected partly liquid and
Fig. 12. Petrogenetic model for the Upper Basin Member rhyolites, from caldera collapse to eruption. M, I and S refer to mafic, intermediate
and silicic replenishing magmas, respectively. Rhyolite unit abbreviations as in Table 1; LCT, Lava Creek Tuff; HRT, Huckleberry Ridge Tuff.
Illustrations are not to scale.
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VOLUME 50
NUMBER 11
NOVEMBER 2009
Fig. 12. Continued.
solidifying unerupted remnants of the Lava Creek Tuff
magma chamber(s) (Fig. 12c).
Continued magma injections beneath the domes
initiated melting of the now low-d18O intracaldera ignimbrite and/or unerupted Lava Creek Tuff magma (Fig. 12d).
Small volumes of silicic replenishing magmas mixed with
these rhyolitic partial melts, increasing the fraction of
melt and leading to the accumulation of several mainly
liquid magma batches (Fig. 12e). The volume of silicic
replenishing magmas was sufficient to modify the composition of the newly formed resident partial melt and increase
its temperature, resulting in melting of resident crystals
and crystallization of higher-temperature mineral aggregates, but insufficient to appreciably modify its isotopic signature, so that these hybrid magmas maintained a lowd18O signature. Once formed, these newly generated
batches of magma erupted rapidly to form the Upper
Basin Member rhyolites at 500 ka (Fig. 12f).
Shortly before 255 ka, new magma injections beneath
the Mallard Lake Resurgent Dome caused renewed
melting of the rhyolitic protolith. These lavas lack clear
evidence for incorporation of a silicic replenishing
2122
GIRARD & STIX
magma component shortly before eruption. The higher
d18O signatures of these younger rhyolites suggest that (1)
they originate from a region distinct from the resurgent
domes that was less affected by hydrothermal alteration,
or (2) the d18O depletion of the protolith beneath the
Mallard Lake resurgent dome was variable. This magmatic episode led to rejuventation of the rhyolitic reservoir
beneath the Mallard Lake Resurgent Dome and
may have been a precursor for the younger Central
Plateau Member magmatism, which began to erupt
80 kyr later.
with pyroxene analyses and Glenna Keating for her help
with whole-rock analyses. We thank Yan Lavalle¤e, Abby
Peterson and Tyler Barton for their assistance in the field.
Fieldwork was conducted under Research and Collecting
Permit YELL-05478; we thank Jake Lowenstern and John
Eichelberger for their permit proposal reviews and the
Yellowstone National Park Research Permit Office for
their advice. Phil Shane and an anonymous reviewer provided useful suggestions that greatly improved the clarity
of the manuscript.
FU N DI NG
S U M M A RY A N D C O N C L U S I O N S
(1) This study has revealed the existence of a new rhyolite
lava flow, previously mapped as part of the Biscuit Basin
lavas, for which we propose the name East Biscuit Basin
lava flow (EBB). This low-SiO2 rhyolite contains no
quartz and no sanidine and is dominated by an assemblage
of sieved plagioclase and plagioclase^pyroxene^Fe^Tioxide aggregates. Based on analyses of samples of similar
geochemistry in this area (Sturchio et al., 1986), it appears
to be a low-d18O rhyolite.
(2) All other Upper Basin Member rhyolites are highSiO2 rhyolites, containing plagioclase, sanidine and
quartz. Sieved plagioclase crystals and aggregates are present in most UBM rhyolites, together with rounded and
embayed quartz crystals.
(3) Despite various degrees of evolution in terms of mineralogy and major element geochemistry observed
throughout the Upper Basin Member rhyolites, variations
in many trace elements such as Rb, Y, Th, U and REE
are limited and appear buffered, suggesting independent
magma batches rather than a single differentiating
reservoir.
(4) Based on the spatial and temporal distribution
of trace element concentrations and ratios in the caldera,
as well as available d18O information, we suggest the
presence of at least six magma batches in early postcaldera time.
(5) These observations lead us to develop a petrogenetic
model in which (a) mafic magma replenishment occurred
beneath two regions of the caldera, triggering uplift of
two resurgent domes; (b) hydrothermal alteration caused
a strong depletion in the d18O signature of the pre-existing
rhyolitic crystal mush and its roof; (c) heat inputs and/or
magma replenishment initiated melting of the low-d18O
mush and/or roof; (d) silicic magma replenishment
enhanced melting and triggered eruptions.
AC K N O W L E D G E M E N T S
We are grafeful to Bill Minarik and Lang Shi for their
invaluable assistance with the ICP-MS and electron
microprobe work, respectively, Crystal Mann for her help
This research was funded through grants to J.S. from the
Natural Sciences and Engineering Research Council of
Canada and le Fonds Que¤be¤cois pour la Recherche sur la
Nature et les Technologies. G.G. acknowledges a scholarship from GEOTOP-Universite¤ du Que¤bec a' Montre¤alMcGill .
S U P P L E M E N TA RY DATA
Supplementary data are available at Journal of Petrology
online.
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Magma Recharge and Crystal Mush Rejuvenation

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