Chemical Geology 236 (2007) 134 – 166
www.elsevier.com/locate/chemgeo
40
Ar/ 39 Ar and U–Pb dating of the Fish Canyon magmatic system,
San Juan Volcanic field, Colorado: Evidence for an extended
crystallization history
O. Bachmann a,⁎, F. Oberli b , M.A. Dungan a , M. Meier b , R. Mundil c , H. Fischer d
a
Section des Sciences de la Terre, Université de Genève, 13, rue des maraîchers, 1211 Genève 4, Switzerland
b
Institute of Isotope Geochemistry and Mineral Resources, ETH Zurich, CH-8092 Zurich, Switzerland
c
Berkeley Geochronology Center, 2455 Ridge Road, Berkeley, CA 94709, USA
d
Uf de Breiti 3, CH-8460 Marthalen, Switzerland
Received 7 October 2004; received in revised form 21 September 2006; accepted 22 September 2006
Editor: P. Deines
Abstract
The ∼ 5000 km3 Fish Canyon Tuff (FCT) is an important unit for the geochronological community because its sanidine, zircon
and apatite are widely used as standards for the 40Ar/39Ar and fission track dating techniques. The recognition, more than 10 years
ago [Oberli, F., Fischer, H. and Meier, M., 1990. High-resolution 238U–206Pb zircon dating of Tertiary bentonites and Fish Canyon
Tuff; a test for age “concordance” by single-crystal analysis. Seventh International Conference on Geochronology,
Cosmochronology and Isotope Geology. Geological Society of Australia Special Publication Canberra, 27:74], of a ≥ 0.4 Ma
age difference between the U–Pb zircon ages and 40Ar/39Ar sanidine ages has, therefore, motivated efforts to resolve the origin of
this discrepancy. To address this controversial issue, we initially performed 37 U–Pb analyses on mainly air-abraded zircons at
ETH Zurich and nearly 200 40Ar/39Ar measurements on hornblende, biotite, plagioclase and sanidine obtained at the University of
Geneva, using samples keyed to a refined eruptive stratigraphy of the FCT magmatic system.
Disequilibrium-corrected 206Pb/238U ages obtained for 29 single-crystal and three multi-grain analyses span an interval of
∼ 28.67–28.03 Ma and yield a weighted mean age of 28.37 ± 0.05 Ma (95% confidence level), with MSWD = 8.4. The individual
dates resolve a range of ages in excess of analytical precision, covering ∼ 600 ka. In order to independently confirm the observed
spread in zircon ages, 12 additional analyses were carried out at the Berkeley Geochronology Center (BGC) on individual zircons
from a single lithological unit, part of them pre-treated by the “chemical abrasion” (CA) technique [Mattinson, J.M., 2005. Zircon
U–Pb chemical abrasion (“CA-TIMS”) method: Combined annealing and multi-step partial dissolution analysis for improved
precision and accuracy of zircon ages. Chemical Geology, 220(1–2): 47–66]. Whereas the bulk of the BGC results displays a
spread overlapping that obtained at ETH, the group of CA treated zircons yield a considerably narrower range with a mean age of
28.61 ± 0.08 Ma (MSWD = 1.0). Both mean zircon ages determined at ETH and BGC are older than the ∼ 28.0 Ma 40Ar/39Ar
eruption age of FCT – even when considering the possibility that the latter may be low by as much as ∼ 1% due to a miscalibration
of the 40K decay constants – and is thus indicative of a substantial time gap between magma crystallization and extrusion. The CA
technique further reveals that younger FCT zircon ages are likely to be associated with chemically unstable U-enriched domains,
which may be linked to crystallization during extended magma residence or may have been affected by pre-eruptive and/or post-
⁎ Corresponding author. Tel.: +41 22 3796893.
E-mail addresses: olivier.bachmann@terre.unige.ch (O. Bachmann), oberli@erdw.ethz.ch (F. Oberli), michael.dungan@terre.unige.ch
(M.A. Dungan), martin.meier@erdw.ethz.ch (M. Meier), rmundil@bgc.org (R. Mundil), hhfischer@access.ch (H. Fischer).
0009-2541/$ - see front matter © 2006 Elsevier B.V. All rights reserved.
doi:10.1016/j.chemgeo.2006.09.005
O. Bachmann et al. / Chemical Geology 236 (2007) 134–166
135
eruptive secondary loss of radiogenic lead. Due to their complex crystallization history and/or age bias due to Pb loss, the FCT
zircon ages are deemed unsuitable for an accurate age calibration of FCT sandine as a fluence monitor for the 40Ar/39Ar method.
Even though data statistics preclude unambiguous conclusions, 40Ar/39Ar dating of sanidine, plagioclase, biotite, and
hornblende from the same sample of vitrophyric Fish Canyon Tuff supports the idea of a protracted crystallization history.
Sanidine, thought to be the mineral with the lowest closure temperature, yielded the youngest age (28.04 ± 0.18 Ma at 95% c.l.,
using Taylor Creek Rhyolite [Renne, P.R. et al., 1998. Intercalibration of standards, absolute ages and uncertainties in 40Ar/39Ar
dating. Chemical Geology, 145: 117–152.] as the fluence monitor), whereas more retentive biotite, hornblende and plagioclase
gave slightly older nominal ages (by 0.2–0.3 Ma). In addition, a laser step-heating experiment on a 2-cm diameter feldspar
megacryst produced a “staircase” argon release spectrum (older ages at higher laser power), suggestive of traces of inherited argon
in the system. Thermal and water budgets for the Fish Canyon magma indicate that the body remained above its solidus (∼ 700 °C)
for an extended period of time (N 105 years). At these temperatures, argon volume diffusion is thought to be fast enough to prevent
accumulation of radiogenic Ar. If this statement were true, an existing isotopic record should have been completely reset within a
few hundred years, regardless of the phase and initial age of the phenocryst. As these minerals are unlikely to be xenocrysts that
were incorporated within such a short time span prior to eruption, we suggest that a fraction of radiogenic Ar can be retained
N 105 years, even at T ∼ 700 °C.
© 2006 Elsevier B.V. All rights reserved.
Keywords: Geochronology; Igneous processes; Magma chambers; Fish Canyon Tuff
1. Introduction
With the wealth of geochronological methods that are
currently available, an increasing number of volcanic
rocks have been dated by several decay schemes on
different minerals, with, at times, perplexing outcomes.
One such unit is the Fish Canyon Tuff, which has been
the target of multiple geochronological studies, in part
due to its size (one of the largest known ignimbrites,
with a volume in excess of 5000 km3), but also because
it is widely used as a natural standard for the 40Ar/39Ar
and fission-track techniques (Cebula et al., 1986). Its
comprehensive geochronological database, with results
from K–Ar, Ar–Ar, Rb–Sr, and U–Pb techniques,
shows a lack of convergence towards a unique, precise
value (total interlaboratory range in age N3.5%; see
Table 1 of Daze et al., 2003; Table 1 of Spell and
McDougall, 2003; Fig. 7 of Schmitz and Bowring,
2001). Particularly important is the fact that the two
most precise methods (40 Ar/ 39 Ar and U–Pb) are
discrepant by ∼ 0.4 Ma. U–Pb ages on zircon and
titanite converge around 28.4–28.5 Ma (Oberli et al.,
1990; Schmitz and Bowring, 2001), whereas 40Ar/39Ar
commonly yields ages around ∼28.0 Ma (Renne et al.,
1994; Renne and Min, 1998; Renne et al., 1998;
Villeneuve et al., 2000). Probable causes for this lack of
convergence towards a unique, precise age include: (1)
calibration problems and/or uncertainties in decay
constants associated with the different dating techniques
(in particular, with the decay constant of 40K (λ);
Lanphere and Dalrymple, 2000; Min et al., 2000;
Schmitz and Bowring, 2001; Kwon et al., 2002;
Schoene et al., 2006) and (2) differences in the apparent
age of the mineral phases present in the magma.
Different elements are known to diffuse at different
rates depending on a number of factors (e.g., Lee, 1995).
Typically, Pb in zircon and titanite diffuses more slowly
than does Ar in feldspar for a given temperature (Foland,
1994; Lee et al., 1997; McDougall and Harrison, 1999;
Cherniak and Watson, 2001), thereby more likely
retaining a memory of pre-eruptive crystallization
episodes even if minerals have had long residence
times at near-solidus temperatures (e.g., Reid et al.,
1997; Brown and Fletcher, 1999; Bacon and Lowenstern, 2005; Charlier et al., 2005).
In an attempt to shed some light on these issues, we
have combined high-precision U–Pb dating of single
zircons by TIMS with 40 Ar/ 39 Ar total-fusion and
incremental-heating experiments on sanidine, biotite,
plagioclase and hornblende from several samples of the
Fish Canyon magmatic system, including co-magmatic
xenoliths found in the intracaldera Fish Canyon Tuff
(Bachmann et al., 2002). The main goals were: (1) to
investigate whether zircons have recorded a single
crystallization event close to the time of eruption, as
suggested by Schmitz and Bowring (2001) or a protracted
period of crystallization (Oberli et al., 1990; Bachmann
et al., 2002, Oberli et al., 2002), and (2) to determine
whether the variable susceptibility of the four main
mineral phases with respect to argon diffusion would
provide evidence for inheritance. In order to simultaneously assess the potential complications arising from the
fact that the Fish Canyon magma erupted in three discrete
events, sanidine and zircon from the three lithologies of
the Fish Canyon magmatic system (the precursory Pagosa
Peak Dacite, the climactic Fish Canyon Tuff, and the post-
136
O. Bachmann et al. / Chemical Geology 236 (2007) 134–166
caldera Nutras Creek Dacite) have also been dated by
40
Ar/39Ar and U–Pb.
2. The Fish Canyon magmatic system
The Fish Canyon magmatic system belongs to the
voluminous mid-Tertiary high-K calc-alkaline ignimbrite
sequence of the San Juan volcanic field in present day
Colorado (Lipman, 2000), and it comprises three separate,
but compositionally identical units that were erupted at
∼28 Ma (Fig. 1). The ∼5000 km3 Fish Canyon Tuff forms
the bulk of the erupted volume (N95%) and was emplaced
as highly mobile ash-flows that covered N10,000 km2
during the collapse of the ∼80 ×30 km La Garita caldera.
The absence of welding breaks in the outflow facies or in
exposures of the intracaldera tuff suggests rapid emplacement (on the order of days?). On the basis of decompression-induced granophyre crystallization as overgrowths on
phenocrysts from the northern intracaldera tuff and the
segmented aspect of the La Garita caldera, the eruption is
thought to have started in the south before propagating
northward (Lipman et al., 1997).
This vast ignimbrite was preceded by eruption of the
precaldera Pagosa Peak Dacite, a poorly fragmented
200 km3 pyroclastic deposit that is distributed around the
southern margin of the La Garita caldera. This unit is
thought to have resulted from low-energy fountaining of
Fish Canyon magma (Bachmann et al., 2000). Although
welding breaks can be observed locally, this unit was also
emplaced rapidly as sections thicker than a few hundreds
of meters remained hot enough to flow rheomorphically
(Bachmann et al., 2000). In rare instances where the
contact between the Pagosa Peak Dacite and the Fish
Canyon Tuff is well exposed, the Fish Canyon Tuff rests
directly on the top of thick, rheomorphic Pagosa Peak
Dacite. Neither erosion, soil formation, nor sediment
deposition took place between the two eruptions,
suggesting a relatively short time gap. However, the
base of the Fish Canyon Tuff is non-welded and does not
show fumarolic alteration, which would be expected if the
Pagosa Peak Dacite had still been hot at the time of the
Fish Canyon Tuff deposition. Following the study of
Riehle et al. (1995), such observations suggest a time gap
at least on the order of months between the two eruptions,
Fig. 1. Simplified location map of the San Juan volcanic field, showing the distribution of the three units of the Fish Canyon magmatic system and the
sampling localities (modified from Bachmann et al., 2002).
O. Bachmann et al. / Chemical Geology 236 (2007) 134–166
although it remained short enough to preclude any
significant erosion or deposition.
A small post-Fish Canyon Tuff lava flow (b1 km3),
the Nutras Creek Dacite, overlies intracaldera tuff on the
northern flank of the resurgent dome of the La Garita
caldera. It is characterized by exposures of devitrified,
flow-banded Fish Canyon magma. The base of the unit
is not exposed and the time gap between this flow and
the Fish Canyon Tuff cannot be assessed on the basis of
field relations.
The Fish Canyon magma displays textural and
geochemical evidence for simultaneous dissolution of
feldspar + quartz and crystallization of hydrous phases
(hornblende + biotite) during gradual near-isobaric
reheating from ∼ 720 to 760 °C (Bachmann and
Dungan, 2002; Bachmann et al., 2002, 2005). This
scenario, along with the high crystallinity (45% crystals)
and near-solidus mineral assemblage of this magma,
suggests that the Fish Canyon magma cooled to a
nearly-solidified, rigid crystal mush before being partly
remelted (“rejuvenated”) by dissolution of feldspar and
quartz prior to eruption. Complete solidification of the
entire system seems unlikely as the thermal and water
inputs necessary to remelt N 50% of the total volume by
dehydration melting are prohibitive (Bachmann et al.,
2002). Nonetheless, co-magmatic holocrystalline xenoliths, which record marginal solidification of the magma
body, are present in the intracaldera Fish Canyon Tuff.
We postulate that this retrograde-prograde temperature path is a consequence of voluminous shallow
intrusions of water-rich mafic magma at the base of the
partly solidified Fish Canyon magma chamber. The
absence of any measurable thermal or chemical
gradients in the Fish Canyon magma cannot be
reconciled with a reheating event dominated by
conductive heat transfer or significant mixing with
mafic magma: it requires that heat was dispersed
throughout the batholithic Fish Canyon chamber rapidly
and pervasively, in order to partially remelt feldspars
and quartz without destabilizing the hydrous phases. On
the basis of numerical simulations, Bachmann and
Bergantz (2003) suggest that upward percolation of a
hot water-rich fluid phase through the Fish Canyon
crystal mush over 150–200 ka is sufficient to account
for the rejuvenation of the magma.
3.
40
Ar/39Ar dating
3.1. Sample preparation
Mineral separates were prepared from two Pagosa
Peak Dacite samples (PPDcc and PPDlc) and from one
137
each of the Fish Canyon Tuff (FCTar) and Nutras Creek
Dacite (NCD; see locations on Fig. 1 and Table 1). The
Nutras Creek Dacite is devitrified, but the samples from
the outflow Fish Canyon Tuff and the Pagosa Peak
Dacite are basal vitrophyres. In view of the abundance
of glassy material in the Pagosa Peak Dacite, two
samples from different localities were dated in order to
evaluate reproducibility. Rock samples were crushed in
a stainless steel mortar and sieved to obtain the optimal
size fractions for the various mineral phases (typically
125–315 μm). Quartz and feldspars were then magnetically separated from the remaining material using a
Frantz magnetic separator. Hornblende and biotite were
hand-picked from the magnetic residue. To obtain
sanidine- and plagioclase-rich fractions, two heavy
liquid steps (using sodium polytungstate) were required.
Sanidine and plagioclase separates were then mildly
etched in dilute HF (2%) for a few minutes and rinsed
with distilled water. All separates were cleaned by
ultrasonic agitation in acetone.
Hand picked aliquots and neutron fluence monitors
(Taylor Creek Rhyolite sanidine) were wrapped in
aluminum foil and placed in wells drilled into 99.99%
pure copper planchettes, before being irradiated for 50
h at the Oregon State University Triga reactor. These
copper planchettes, with small wells for neutron fluence
monitors surrounding the larger ones containing the
unknowns, were designed to minimize the error on the J
parameter in placing the monitors physically as close as
Table 1
Coordinates of sample location
Sample # Location
(quadrangle)
Unit
PPDcc
PPD
Columbine Creek
(Mt Hope)
PPDlc
Lake Creek (Wolf
Creek Pass)
FCTar
Agua Ramon
(South Fork East)
FCTfv
Fun Valley (Beaver
Creek Reservoir)
NCD
Nutras Creek
(Elk Park)
MegaX
Willow Creek
(San Luis Peak)
TonX
Willow Creek
(San Luis Peak)
GrdX1,
Machin Lake area
GrdX2 (Halfmoon Pass)
GrnX
Machin Lake area
(Halfmoon Pass)
Lat.
Long.
37°34′15″ 106°45′
30″
PPD
37°29′17″ 106°52′
12″
FCT
37°42′40″ 106°33′
(outflow)
22″
FCT
37°36′48″ 106°42′
(outflow)
08″
NCD
38°01′37″ 106°50′
03″
FCT
37°55′39″ 106°53′
(intracaldera)
46″
Tonalitic
37°55′39″ 106°53′
xenolith
46″
Granodioritic 37°56′14″ 106°51′
xenoliths
32″
Granitic
37°55′35″ 106°51′
xenolith
00″
PPD = Pagosa Peak Dacite, FCT = Fish Canyon Tuff, NCD = Nutras
Creek Dacite, MegaX = Megacryst.
138
O. Bachmann et al. / Chemical Geology 236 (2007) 134–166
possible to the unknowns. A J value was calculated for
each sample using the three wells containing the fluence
monitors that surround the sample well. The uncertainty
for the J parameter was found to be reproducible and
typically ∼ 0.3% of the age (1σ).
3.2.
40
40
Ar/39Ar analyses
Ar/39 Ar analyses were performed at the University
of Geneva. For details on equipment and analytical
procedures, the reader is referred to Singer et al. (1999).
1σ errors of individual analyses (mentioned in text and
listed in Table 2) include precision of isotopic ratio
measurements, reproducibility of the system blank and
mass discrimination of the spectrometer, as well as
uncertainties associated with correction for interfering
reactions during irradiation. These errors were obtained
by quadratic summation and were applied to the calculation of inverse-variance weighted mean ages and
isochron parameters, their standard deviations, and
MSWD values. The uncertainty on J, which typically
dominates the total error estimate, was then propagated
into the final error of mean, plateau, and isochron ages.
Ages were calculated with respect to an age of 28.34 Ma
for the Taylor Creek Rhyolite (Renne et al., 1998) and
are based on decay constants recommended by Steiger
and Jäger (1977).
Both total-fusion and incremental-heating analyses
were performed using a CO2 laser. Total-fusion experiments allowed the acquisition of high precision singlegrain data on biotite and sanidine at the chosen size
fraction (125–315 μm), but required aliquots of 10 to 15
grains for K-poor minerals (plagioclase and hornblende). Incremental-heating analyses, achieved by
stepwise increase in laser power, were undertaken on
all four mineral phases (sanidine, biotite, plagioclase,
hornblende), as well as on a ∼ 2 cm feldspar megacryst
extracted from an intracaldera Fish Canyon Tuff
pumice. Experiments on K-rich minerals (sanidine and
biotite) required aliquots of 10 to 15 grains, whereas
larger quantities (up to 50 grains) were needed for
plagioclase and hornblende measurements. Depending
on sample size and degassing behavior, the number of
steps varied from less than 10 to around 20, except for
the megacryst, which was large enough to permit a 24step experiment. Calculation and assessment of the
precision of plateau and isochron ages determined from
incremental analyses are based on the recommendations
of McDougall and Harrison (1999) and Singer et al.
(1999). Except for the FCTar sanidine and biotite
experiments, a few steps (those significantly increasing
MSWD) were omitted from the plateau calculations.
However, plateau ages were calculated from more than
94% of the gas released for all mineral phases.
3.3. Results
3.3.1. Sanidine
Due to its high potassium concentration and status as
an international neutron fluence monitor, Fish Canyon
sanidine (appropriate disordered structural state confirmed by XRD measurements; Whitney and Stormer,
1985) was dated from all three units of the Fish Canyon
system by both total-fusion and incremental-heating
methods. The single-grain total-fusion results for the
three lithological units are reproducible (Table 2); out of
more than 60 analyses performed, only one gave an
aberrant age of 25.78 ± 0.09 Ma (in PPDcc, Table 2) and
was not included in the calculations. The two Pagosa
Peak Dacite samples yielded identical weighted mean ages
of 27.93 ± 0.09 Ma (MSWD= 2.3) and 27.94 ± 0.09 Ma
(MSWD = 1.0). Similarly, the Fish Canyon Tuff and
Nutras Creek Dacite gave indistinguishable ages of
28.04 ± 0.09 Ma (MSWD = 0.2) and 28.07 ± 0.09 Ma
(MSWD = 1.0).
All incremental-heating experiments produced welldefined plateaus comprising more than 98% of the gas
(Tables 4 and 5; Fig. 2). A few gas release steps at very
low laser power were omitted from the age calculations.
These low temperature steps yielded aberrant ages probably due to incomplete radiogenic argon retention near
the grain surfaces (e.g., Albarède, 1978). Weighted mean
ages of 27.94 ± 0.09 Ma (MSWD = 2.3) and 28.04 ±
0.09 Ma (MSWD = 2.0) were obtained for the Pagosa
Peak Dacite and Fish Canyon Tuff, respectively. Two
replicates of the Nutras Creek Dacite gave 28.08 ± 0.09
(MSWD = 2.6) and 28.05 ± 0.09 Ma (MSWD = 0.5).
Inverse isochron ages calculated for the incrementalheating analyses are consistent with plateau and totalfusion results (PPDcc = 27.96 ± 0.09 Ma, MSDW = 2.4;
FCTar = 28.01 ± 0.09 Ma, MSDW = 1.1; NCD = 28.02 ±
0.09 and 28.06 ± 0.09 Ma, MSWD = 0.8, 0.5). Two out of
four 36Ar/40Ar intercepts are not atmospheric within error
(FCTar and NCD#1 sanidines), but this is most likely due
to the extremely high fractions of radiogenic argon
yielded by all steps, giving imprecise regression lines.
Apart from these two analyses, all other 36Ar/40Ar
intercepts are atmospheric within errors, including the
other NCD sanidine (#2; from the same sample; Fig. 2)
and a biotite analysis on FCTar (Fig. 3).
The ages obtained for all three units by both totalfusion and step-heating of sanidine are indistinguishable
at the 1σ level. It is noted, however, that the ages obtained
for both samples of Pagosa Peak Dacite are nominally
O. Bachmann et al. / Chemical Geology 236 (2007) 134–166
139
Table 2
Ar/39Ar isotopic data for individual analyses
40
40
Ar/39Ara
Total fusion analyses
FCTar (san)
J = 0.0135981
GE111C1A
1.143632
GE111C1B
1.137895
GE111C1Q
1.141411
GE111C1R
1.142623
GE111C1S
1.142095
GE111C1T
1.141244
GE111C1U
1.141469
GE111C1V
1.141990
GE111C1W
1.140724
GE111C1X
1.143119
GE111C1Y
1.144296
GE111C1Z
1.140346
FCTar (plag)
J = 0.0134704
GE112C2A
1.188852
GE112C2B
1.170925
GE112C2C
0.185213
FCTar (hbl)
J = 0.0134704
GE113C1A
1.643527
GE113C1B
1.741333
GE113C1C
1.575396
FCTar (bio)
J = 0.013413
GE112C4A
1.304446
GE112C4B
1.288435
GE112C4C
1.455079
GE112C4D
1.475214
PPDlc (Bio)
J = 0.01341055
GE112C3A
1.299564
GE112C3B
1.330194
GE112C3C
1.340554
GE112C3D
1.277122
PPDlc (san)
J = 0.01365357
GE111C3A
1.130576
GE111C3B
1.137321
GE111C3C
1.130359
GE111C3D
1.133289
GE111C3E
1.137308
GE111C3F
1.130172
GE111C3G
1.130322
J = 0.01361595
NCD (san)
GE111C2A
1.142589
GE111C2B
1.153482
GE111C2C
1.140593
GE111C2D
1.158803
GE111C2E
1.146754
GE111C2F
1.143278
GE111C2G
1.140944
GE111C2H
1.143054
GE111C2I
1.142927
GE111C2J
1.140710
GE111C2K
1.144080
GE111C2L
1.143661
GE111C2M
1.143490
GE111C2N
1.143596
GE111C2O
1.140572
GE111C2P
1.143522
37
Ar/39Ara
N = 12
0.007849
0.006783
0.006640
0.006735
0.006690
0.007216
0.006979
0.006738
0.006783
0.006459
0.007573
0.006538
N=3
2.965818
2.949725
2.981663
N=3
6.476819
6.439619
6.430503
N=4
0.010261
0.031729
0.012159
0.011877
N=4
0.021083
0.022642
0.012249
0.012947
N=7
0.006689
0.006794
0.006969
0.007397
0.007025
0.006975
0.006513
N = 27
0.007385
0.007113
0.007079
0.007967
0.006716
0.006434
0.007606
0.188154
0.006775
0.006435
0.006811
0.006973
0.006322
0.006661
0.006771
0.006683
36
40
Ar⁎
(10−14 mol)
%40Ar⁎
K/Ca
Apparent ageb
(Ma) ± 1σ
0.00000947
0.00000250
0.00001290
0.00001130
0.00001300
0.00001300
0.00000881
0.00001080
0.00001087
0.00001429
0.00001932
0.00000044
3.170645
3.921292
1.370600
1.615103
2.103951
1.455636
1.501582
1.464530
1.411640
1.652432
1.873154
2.033967
99.40
99.57
99.30
99.35
99.30
99.31
99.41
99.36
99.36
99.27
99.15
99.63
62.4
72.2
73.8
72.8
73.2
67.9
70.2
72.7
72.2
75.9
64.7
74.9
28.09 ± 0.05
28.00 ± 0.07
28.01 ± 0.05
28.05 ± 0.05
28.02 ± 0.07
28.00 ± 0.05
28.04 ± 0.05
28.04 ± 0.05
28.01 ± 0.07
28.04 ± 0.05
28.03 ± 0.05
28.07 ± 0.07
⁎
⁎
⁎
⁎
⁎
⁎
⁎
⁎
⁎
⁎
⁎
⁎
0.00091171
0.00082255
0.00088177
0.506163
0.636561
0.405925
96.70
98.80
97.55
0.2
0.2
0.2
28.20 ± 0.12
28.37 ± 0.11
28.36 ± 0.14
⁎
⁎
⁎
0.00332080
0.00360046
0.00308951
0.750628
0.896456
0.819715
71.22
66.91
74.08
7.5
7.6
0.1
28.17 ± 0.16
28.30 ± 0.16
28.22 ± 0.14
⁎
⁎
⁎
0.00048536
0.00041952
0.00100132
0.00105205
0.913047
1.103549
0.383163
0.477969
88.71
90.21
79.41
78.67
47.8
15.4
40.3
41.3
28.20 ± 0.07
28.33 ± 0.07
28.16 ± 0.11
28.29 ± 0.11
⁎
⁎
⁎
⁎
0.00046387
0.00057502
0.00060649
0.00037412
0.317999
0.755392
0.557696
0.589160
89.22
87.01
86.36
91.06
23.2
21.6
40.0
37.9
28.25 ± 0.10
28.20 ± 0.10
28.21 ± 0.10
28.34 ± 0.10
⁎
⁎
⁎
⁎
0.00001074
0.00002190
0.00000725
0.00000252
0.00000970
0.00000527
0.00000700
1.460864
1.878863
4.395066
0.877185
0.890388
1.647937
1.279671
99.35
99.07
99.45
99.58
99.39
99.50
99.45
73.3
72.1
70.3
66.3
69.8
70.3
75.2
27.87 ± 0.07
27.95 ± 0.09
27.89 ± 0.07
28.00 ± 0.11
28.04 ± 0.05
27.90 ± 0.07
27.89 ± 0.07
⁎
⁎
⁎
⁎
⁎
⁎
⁎
0.00000885
0.00004993
0.00000575
0.00006682
0.00001389
0.00001126
0.00000910
0.00005922
0.00001238
0.00001076
0.00002482
0.00001451
0.00001767
0.00001409
0.00001355
0.00002516
3.706494
1.800781
1.042908
0.830779
1.461081
1.759253
1.254725
0.907363
1.293246
1.381541
2.166046
1.614954
1.307181
1.516003
1.512893
0.951612
99.42
98.37
99.49
97.95
99.28
99.35
99.41
99.37
99.32
99.36
99.00
99.27
99.18
99.28
99.29
98.99
66.4
68.9
69.2
61.5
73.0
76.2
64.4
2.6
72.3
76.1
71.9
70.3
77.5
73.6
72.4
73.3
28.10 ± 0.10
28.07 ± 0.12
28.08 ± 0.07
28.08 ± 0.07
28.17 ± 0.07
28.10 ± 0.05
28.06 ± 0.07
28.10 ± 0.07
28.08 ± 0.07
28.04 ± 0.05
28.02 ± 0.07
28.09 ± 0.07
28.06 ± 0.07
28.09 ± 0.05
28.02 ± 0.07
28.01 ± 0.07
⁎
⁎
⁎
⁎
⁎
⁎
⁎
⁎
⁎
⁎
⁎
⁎
⁎
⁎
⁎
⁎
Ar/39Ara
Step used in regression
(continued on next page)
140
O. Bachmann et al. / Chemical Geology 236 (2007) 134–166
Table 2 (continued )
40
Ar/39Ara
37
Ar/39Ara
Total fusion analyses
NCD (san)
J = 0.01361595
GE111C2Q
1.141447
GE111C2R
1.145087
GE111C2S
1.142035
GE111C2T
1.141914
GE111C2U
1.141414
GE111C2V
1.142007
GE111C2W
1.141650
GE111C2X
1.141637
GE111C2Y
1.142860
GE111C2Z
1.143630
GE111C21
1.143112
PPDcc (san)
J = 0.01329567
GE111C4A
1.132945
GE111C4B
1.170666
GE111C4C
1.134965
GE111C4D
1.131631
GE111C4E
1.143077
GE111C4F
1.130399
GE111C4G
1.131810
GE111C4H
1.138108
GE111C4I
1.137017
GE111C4J
1.126837
GE111C4K
1.130749
GE111C4L
1.133083
GE111C4M
1.155324
GE111C4N
1.137804
GE111C4O
1.128812
GE111C4P
1.133724
N = 27
0.006884
0.007630
0.006998
0.006765
0.007192
0.006583
0.007043
0.006521
0.006121
0.007092
0.006438
N = 16
0.007680
0.010949
0.007635
0.007382
0.006662
0.009385
0.007315
0.007896
0.007374
0.007554
0.007334
0.008341
0.008003
0.006869
0.006954
0.006821
Incremental heating analyses
FCTar (san)
J = 0.0135981
GE111C1C
1.515267
GE111C1E
1.155765
GE111C1F
1.143149
GE111C1G
1.140620
GE111C1H
1.139337
GE111C1I
1.137337
GE111C1K
1.139136
GE111C1L
1.144715
GE111C1M
1.143851
GE111C1N
1.143719
GE111C1O
1.147058
GE111C1P
1.205849
FCTar (hbl)
J = 0.01324858
GE113C1D
25.214670
GE113C1E
6.749717
GE113C1F
9.029773
GE113C1G
3.903025
GE113C1H
2.289285
GE113C1I
1.444540
GE113C1J
1.410654
GE113C1K
1.325337
GE113C1L
1.269737
GE113C1M
1.351698
GE113C1N
1.685116
N = 12
0.016252
0.001075
0.008958
0.007994
0.007179
0.006685
0.007674
0.007680
0.006491
0.006513
0.006731
0.007133
N = 11
0.234863
0.204102
0.404456
1.728248
5.558185
5.912695
5.968434
6.041674
5.173800
6.681218
15.662040
36
40
Ar⁎
(10−14 mol)
%40Ar⁎
K/Ca
Apparent ageb
(Ma) ± 1σ
0.00000810
0.00001342
0.00001889
0.00001745
0.00001427
0.00001891
0.00001030
0.00000720
0.00001110
0.00002161
0.00001586
1.845976
1.478376
1.401680
0.776706
1.373543
1.364709
1.598307
0.942999
1.992892
1.841639
1.517976
99.43
99.30
99.15
99.19
99.27
99.15
99.37
99.45
99.35
99.08
99.23
71.2
64.2
70.0
72.4
68.1
74.4
69.6
75.1
80.1
69.1
76.1
28.08 ± 0.05
28.13 ± 0.07
28.02 ± 0.07
28.02 ± 0.07
28.03 ± 0.07
28.01 ± 0.07
28.07 ± 0.05
28.09 ± 0.07
28.09 ± 0.07
28.04 ± 0.07
28.06 ± 0.07
⁎
⁎
⁎
⁎
⁎
⁎
⁎
⁎
⁎
⁎
⁎
0.00003691
0.00044013
0.00001977
0.00001305
0.00002104
0.00001680
0.00001301
0.00002220
0.00003279
0.00001502
0.00001219
0.00003304
0.00008511
0.00001636
0.00001845
0.00001524
3.543552
3.786444
1.340695
1.192718
2.070621
1.515276
1.955547
1.360739
0.945939
1.815113
1.552954
1.973311
1.794297
1.611072
2.185201
1.672683
98.69
88.57
99.13
99.30
99.10
99.22
99.30
99.07
98.79
99.25
99.32
98.79
97.48
99.21
99.15
99.24
63.8
44.8
64.2
66.4
73.6
52.2
67.0
62.1
66.5
64.9
66.8
58.8
61.2
71.3
70.5
71.8
27.94 ± 0.07
25.78 ± 0.09
27.95 ± 0.06
27.92 ± 0.07
28.14 ± 0.06
27.87 ± 0.07
27.92 ± 0.06
28.01 ± 0.07
27.91 ± 0.07
27.79 ± 0.06
27.90 ± 0.07
27.81 ± 0.05
27.98 ± 0.07
28.05 ± 0.07
27.81 ± 0.07
27.95 ± 0.07
⁎
0.00112693
0.00003956
0.00003277
0.00001956
0.00000964
0.00000591
0.00000556
0.00002466
0.00002406
0.00000128
0.00001734
0.00019532
0.055119
0.330056
0.381530
0.726393
0.919791
2.175118
2.107747
0.372415
0.360246
1.738653
1.432190
0.577138
77.80
98.66
98.81
99.14
99.39
99.48
99.49
99.00
99.02
99.60
99.19
94.87
30.2
45.6
54.7
61.3
68.3
73.3
72.7
72.1
75.5
75.2
72.8
68.7
29.12 ± 0.77
28.17 ± 0.15
27.91 ± 0.11
27.93 ± 0.09
27.98 ± 0.05
27.96 ± 0.05
28.00 ± 0.07
28.00 ± 0.09
27.99 ± 0.10
28.15 ± 0.05
28.11 ± 0.07
28.27 ± 0.08
⁎
⁎
⁎
⁎
⁎
⁎
⁎
⁎
⁎
⁎
⁎
⁎
0.08256177
0.01889007
0.02681819
0.00967113
0.00532586
0.00256052
0.00251420
0.00232870
0.00205944
0.00232554
0.00572722
0.009062
0.015517
0.047635
0.037028
0.469746
0.550023
0.597246
0.386733
0.193305
0.070397
0.007190
3.29
17.42
12.54
30.17
50.28
79.72
80.51
83.83
83.98
87.96
72.90
2.1
2.4
1.2
0.3
0.1
0.1
0.1
0.1
0.1
0.1
0.0
20.07 ± 4.19
28.31 ± 3.00
27.27 ± 2.02
28.38 ± 1.04
27.82 ± 0.19
27.84 ± 0.14
27.46 ± 0.15
26.87 ± 0.23
25.78 ± 0.32
28.75 ± 0.92
29.89 ± 0.96
⁎
⁎
⁎
⁎
⁎
⁎
⁎
⁎
⁎
Ar/39Ara
Step used in regression
⁎
⁎
⁎
⁎
⁎
⁎
⁎
⁎
⁎
⁎
⁎
⁎
⁎
⁎
O. Bachmann et al. / Chemical Geology 236 (2007) 134–166
141
Table 2 (continued )
40
Ar/39Ara
Incremental heating analyses
FCTar (bio)
J = 0.013413
GE112C4E
4.632346
GE112C4F
2.059564
GE112C4G
1.543156
GE112C4H
1.331533
GE112C4I
1.323980
GE112C4J
1.307236
GE112C4K
1.282739
GE112C4L
1.295064
GE112C4M
1.262225
GE112C4N
1.250063
GE112C4O
1.228021
GE112C4P
1.216517
GE112C4Q
1.197908
FCTar (plag)
J = 0.0134704
GE112C2D
12.779330
GE112C2E
3.655244
GE112C2F
2.176193
GE112C2G
1.227057
GE112C2H
1.167590
GE112C2I
1.159338
GE112C2J
1.162990
GE112C2K
1.168949
GE112C2L
1.154430
GE112C2M
1.153823
GE112C2N
1.156198
GE112C2O
1.355119
PPDlc (bio)
J = 0.01341055
GE112C3E
12.274790
GE112C3F
14.426170
GE112C3G
3.779988
GE112C3H
3.406739
GE112C3I
2.180084
GE112C3J
1.978369
GE112C3K
1.954764
GE112C3L
1.732329
GE112C3M
1.620220
GE112C3N
1.416182
GE112C3O
1.331065
GE112C3P
1.215875
GE112C3Q
1.256295
GE112C3R
16.054910
NCD #1 (san)
J = 0.01361595
GE111214
1.286336
GE111215
1.317837
GE111216
1.297691
GE111217
1.226535
GE111218
1.179669
GE111219
1.144946
GE111219
1.140643
GE111219
1.138790
GE111219
1.137817
GE111219
1.140259
GE111219
1.143943
NCD#2 (san)
J = 0.01361595
GE111C22
3.394290
GE111C23
1.238055
37
Ar/39Ara
N = 13
0.018140
0.009771
0.007962
0.007297
0.007157
0.012144
0.015173
0.027547
0.042203
0.059624
0.046771
0.048496
0.024193
N = 12
2.028861
2.149893
2.531674
2.872734
3.057303
3.092190
3.103459
2.900963
2.982237
3.185253
3.111044
2.998234
N = 14
0.164141
0.075116
0.022265
0.048845
0.025176
0.009949
0.010120
0.010783
0.009132
0.011426
0.039176
0.029371
0.008345
0.062058
N = 11
0.124377
0.012409
0.016424
0.012201
0.009644
0.008396
0.007356
0.007345
0.006787
0.006592
0.006469
N = 11
0.015334
0.009164
36
40
Ar⁎
(10−14 mol)
%40Ar⁎
K/Ca
Apparent ageb
(Ma) ± 1σ
0.01196300
0.00300200
0.00124416
0.00052428
0.00054607
0.00047232
0.00042883
0.00046516
0.00035941
0.00032212
0.00024585
0.00022561
0.00013423
0.040126
0.090056
0.246674
0.163617
0.314616
0.430935
0.300962
0.471240
0.652765
0.586804
0.467870
0.315912
0.403137
23.62
56.74
75.92
88.06
87.50
89.04
89.85
89.20
91.48
92.39
94.01
94.45
96.46
27.0
50.2
61.5
67.2
68.5
40.4
32.3
17.8
11.6
8.2
10.5
10.1
20.3
26.68 ± 0.98
28.48 ± 0.48
28.55 ± 0.16
28.57 ± 0.20
28.23 ± 0.11
28.37 ± 0.14
28.09 ± 0.12
28.15 ± 0.09
28.14 ± 0.11
28.15 ± 0.09
28.14 ± 0.11
28.01 ± 0.14
28.16 ± 0.12
⁎
⁎
⁎
⁎
⁎
⁎
⁎
⁎
⁎
⁎
⁎
⁎
⁎
0.03846334
0.00932203
0.00409048
0.00106415
0.00089056
0.00087153
0.00084833
0.00078877
0.00079828
0.00085910
0.00080497
0.00099663
0.001063
0.008290
0.022786
0.059374
0.178738
0.262096
0.355802
0.385904
0.453733
0.474187
0.229205
0.169437
12.28
29.17
53.46
92.54
97.80
98.51
99.18
99.32
99.62
99.46
99.33
95.45
0.2
0.2
0.2
0.2
0.2
0.2
0.2
0.2
0.2
0.2
0.2
0.2
38.36 ± 36.76
26.15 ± 5.34
28.52 ± 1.80
27.85 ± 0.48
28.01 ± 0.19
28.01 ± 0.20
28.29 ± 0.16
28.47 ± 0.11
28.21 ± 0.11
28.15 ± 0.15
28.46 ± 0.21
31.69 ± 0.31
⁎
⁎
⁎
⁎
⁎
⁎
⁎
⁎
⁎
⁎
⁎
0.02238711
0.04640748
0.00266025
0.00738257
0.00266371
0.00321495
0.00248500
0.00204589
0.00145163
0.00093580
0.00059550
0.00019283
0.00040381
0.05330003
0.001937
0.000980
0.002723
0.003478
0.012306
0.008168
0.013343
0.016317
0.038303
0.059778
0.292263
0.674640
0.067473
0.001475
46.17
4.95
79.13
35.94
63.77
51.78
62.24
64.88
73.28
80.21
86.66
95.12
90.18
1.90
3.0
6.5
22.0
10.0
19.5
49.3
48.4
45.4
53.7
42.9
12.5
16.7
58.7
7.9
134.15 ± 93.30
17.45 ± 26.08
72.01 ± 35.80
29.82 ± 12.00
33.83 ± 3.96
24.99 ± 4.51
29.63 ± 3.58
27.39 ± 2.68
28.93 ± 1.09
27.68 ± 0.77
28.38 ± 0.15
28.18 ± 0.09
27.61 ± 0.63
7.47 ± 19.56
0.00460755
0.00625179
0.00143507
0.00117044
0.00016495
0.00005560
0.00001652
0.00000720
0.00000350
0.00000275
0.00000982
0.000071
0.000765
0.003968
0.010990
0.081848
0.149604
0.245638
0.738067
2.345597
0.766621
2.299892
5.44
40.46
67.06
71.49
95.54
98.22
99.22
99.46
99.55
99.57
99.38
3.9
39.5
29.8
43.7
50.8
58.4
66.6
66.7
72.2
74.3
75.8
− 1.75 ± 38.49
− 13.34 ± 27.28
21.57 ± 7.44
21.73 ± 3.12
27.89 ± 0.45
27.82 ± 0.27
28.00 ± 0.17
28.02 ± 0.07
28.02 ± 0.05
28.09 ± 0.07
28.13 ± 0.07
⁎
⁎
⁎
⁎
⁎
⁎
⁎
0.00595435
0.00019836
0.005551
0.135273
48.06
94.95
32.0
53.5
40.23 ± 6.46
29.08 ± 0.27
⁎
Ar/39Ara
Step used in regression
⁎
⁎
⁎
⁎
⁎
⁎
(continued on next page)
142
O. Bachmann et al. / Chemical Geology 236 (2007) 134–166
Table 2 (continued )
40
Ar/39Ara
Incremental heating analyses
NCD#2 (san)
J = 0.01361595
GE111C24
1.139121
GE111C25
1.136753
GE111C26
1.136391
GE111C27
1.139532
GE111C28
1.145006
GE111C29
1.145675
GE111C2Z
1.145567
GE111C12
1.147239
GE111C13
1.168215
MegaX
J = 0.01432
GE112A1A
19.895660
GE112A1B
4.048245
GE112A1C
1.696408
GE112A1D
1.379510
GE112A1E
1.393251
GE112A1F
1.154371
GE112A1G
1.157305
GE112A1H
1.156125
GE112A1I
1.181544
GE112A1J
1.219832
GE112A1K
1.204572
GE112A1L
1.192871
GE112A1M
1.184237
GE112A1N
1.233375
GE112A1O
1.240007
GE112A1P
1.225832
GE112A1Q
1.199371
GE112A1R
1.291385
GE112A1S
1.225796
GE112A1T
1.209698
GE112A1U
1.206592
GE112A1V
1.231002
GE112A1W
1.207817
GE112A1X
1.196805
PPDcc (san)
J = 0.01329567
GE111C4Q
1.488483
GE111C4R
1.146072
GE111C4S
1.134260
GE111C4T
1.129736
GE111C4U
1.139961
GE111C4V
1.128413
GE111C4W
1.131792
GE111C4X
1.130612
GE111C4Y
1.146597
GE111C4Z
1.139796
GE111C41
1.140167
GE111C42
1.138577
GE111C43
1.156308
37
Ar/39Ara
N = 11
0.007575
0.007186
0.006915
0.006843
0.006893
0.006713
0.006699
0.006690
0.006966
N = 24
0.011835
0.014239
0.011992
0.011290
0.013775
0.014072
0.015665
0.019165
0.021967
0.023804
0.017584
0.015081
0.012812
0.010626
0.011541
0.010244
0.009842
0.009926
0.007949
0.007060
0.006867
0.007882
0.007228
0.007377
N = 13
0.015367
0.012015
0.012344
0.008805
0.007559
0.007042
0.006970
0.006966
0.007030
0.007288
0.006895
0.007199
0.007031
36
40
Ar⁎
(10−14 mol)
%40Ar⁎
K/Ca
Apparent ageb
(Ma) ± 1σ
0.00001342
0.00000180
0.00000175
0.00000148
0.00000784
0.00000450
0.00000369
0.00001299
0.00008983
1.386110
0.805742
2.337171
1.503773
0.284465
0.868654
0.605347
0.926897
0.123843
99.30
99.59
99.59
99.60
99.44
99.52
99.55
99.31
97.38
64.7
68.2
70.9
71.6
71.1
73.0
73.2
73.3
70.3
27.99 ± 0.07
28.01 ± 0.07
28.00 ± 0.05
28.08 ± 0.09
28.17 ± 0.16
28.21 ± 0.09
28.21 ± 0.24
28.19 ± 0.09
28.14 ± 0.28
⁎
⁎
⁎
⁎
⁎
⁎
⁎
⁎
⁎
0.06372681
0.01010405
0.00221141
0.00104740
0.00105089
0.00022112
0.00015665
0.00020776
0.00033467
0.00045188
0.00038639
0.00035355
0.00031748
0.00048898
0.00052046
0.00046292
0.00034812
0.00064059
0.00043193
0.00033429
0.00038554
0.00043246
0.00032983
0.00029032
0.015087
0.053807
0.062468
0.084631
0.253464
0.125421
0.088227
0.198480
0.297724
0.546601
0.637989
2.154725
1.458139
1.105861
0.472477
0.600263
0.397658
0.210390
0.324657
0.213052
0.300223
0.296057
0.222594
0.254366
5.33
26.16
61.26
77.29
77.46
94.03
95.70
94.42
91.38
88.83
90.25
90.95
91.77
87.98
87.30
88.53
91.10
85.04
89.26
91.50
90.22
89.29
91.59
92.49
41.4
34.4
40.9
43.4
35.6
34.8
32.1
25.6
22.3
20.6
27.9
32.5
38.3
46.1
42.5
47.8
49.8
49.4
61.6
69.4
71.4
62.2
67.8
66.4
27.60 ± 3.96
27.56 ± 1.22
27.05 ± 0.71
27.75 ± 0.54
28.08 ± 0.25
28.24 ± 0.33
28.81 ± 0.49
28.40 ± 0.25
28.10 ± 0.15
28.19 ± 0.10
28.29 ± 0.08
28.23 ± 0.07
28.28 ± 0.07
28.23 ± 0.09
28.17 ± 0.12
28.24 ± 0.10
28.43 ± 0.14
28.57 ± 0.23
28.47 ± 0.14
28.79 ± 0.19
28.32 ± 0.15
28.60 ± 0.16
28.78 ± 0.18
28.80 ± 0.13
⁎
⁎
⁎
⁎
⁎
⁎
⁎
⁎
0.00111217
0.00005893
0.00005259
0.00002254
0.00005278
0.00000718
0.00000084
0.00000631
0.00004812
0.00001070
0.00002293
0.00002392
0.00008891
0.101360
0.205951
0.159221
1.864145
2.021732
2.106350
0.401575
0.349381
0.358633
0.732127
1.113733
0.632368
0.322642
77.69
98.17
98.31
99.06
98.28
99.45
99.62
99.47
98.40
99.37
99.05
99.02
97.37
31.9
36.5
39.7
55.7
64.8
69.6
70.3
70.3
69.7
67.2
71.1
68.1
69.7
28.73 ± 0.35
27.95 ± 0.15
27.71 ± 0.15
27.81 ± 0.06
27.84 ± 0.05
27.88 ± 0.07
28.01 ± 0.09
27.94 ± 0.11
28.03 ± 0.09
28.14 ± 0.07
28.06 ± 0.07
28.01 ± 0.08
27.97 ± 0.12
⁎
⁎
⁎
⁎
⁎
⁎
⁎
⁎
⁎
⁎
⁎
⁎
Ar/39Ara
Step used in regression
Samples were irradiated for 50 h at OSU Triga reactor. Analyses used a CO2 laser and MAP 216 spectrometer at the University of Geneva. Procedures
given in Singer et al. (1999).aCorrected for 37Ar and 39Ar decay: half-lives of 35 days and 259 years respectively. bAll ages are calculated relative to
28.34 Ma Taylor Creek sanidine (Renne et al., 1998) and errors do not include the error on the J-value. Decay constants: λE = 0.581 × 10− 10/year;
λB = 4.962 × 10− 10/year. Power of CO2 laser used: 25 W.
O. Bachmann et al. / Chemical Geology 236 (2007) 134–166
143
Table 3
Summary of total fusion 40Ar/39Ar results on samples from the Fish Canyon magmatic system
Sample #
Material
K2O (wt.%)
K/Ca
a
N
Total fusion age
39
Age (Ma) ± 1σ
MSWD
100
100
100
100
100
86.6
100
100
28.07 ± 0.09
28.04 ± 0.09
28.25 ± 0.09
28.31 ± 0.11
28.23 ± 0.12
27.93 ± 0.09
27.94 ± 0.09
28.25 ± 0.10
1.0
0.2
0.9
0.6
0.2
2.3
1.0
0.4
Ar (%)
NCD
FCTar
FCTar
FCTar
FCTar
PPDcc
PPDlc
PPDlc
Sanidine
Sanidine
Biotite
Plagioclase
Hornblende
Sanidine
Sanidine
Biotite
11
11
9
1
1
11
11
9
69.8
70.7
33.3
0.2
0.2
62.7
71.1
30.8
27 of 27
12 of 12
4 of 4
3 of 3
3 of 3
15 of 16
7 of 7
4 of 4
a
N = number of total fusion analyses used to calculate the mean age. Single grains were used for K-rich minerals (sanidine and biotite), whereas
aliquots of 10-15 grains were used for K-poor minerals (plagioclase and hornblende). K/Ca is the average value for the given number of analyses. See
legend in Table 2 for details on analytical procedures.
younger than those obtained for the overlying Fish
Canyon Tuff and Nutras Creek Dacite. This must be an
artifact lacking chronological significance, as stratigraphic relations unambiguously show that the Pagosa Peak
Dacite was erupted prior to the Fish Canyon Tuff.
3.3.2. Plagioclase, biotite, hornblende
Plagioclase, biotite, and hornblende analyses were
also performed by both total-fusion and step-heating
methods, with the focus on the Fish Canyon Tuff. The
underlying strategy was to assess the potential age
variability between different mineral phases of the same
unit. Biotite from a Pagosa Peak Dacite sample was also
measured in order to compare biotite age variability between samples. As was the case for sanidine, total-fusion
analyses were highly reproducible and all analyses were
included in the calculations. The three phases from the
Fish Canyon Tuff and the biotite from the Pagosa Peak
Dacite gave similar results ranging from 28.23 ± 0.12 Ma
(MSWD = 0.2) for hornblende to 28.31 ± 0.11 Ma
(MSWD = 0.6) for plagioclase (Table 3).
In laser step-heating experiments, biotite yielded
weighted mean plateau ages of 28.22 ± 0.11 Ma
(MSWD = 0.7) for Pagosa Peak Dacite and 28.19 ±
0.09 Ma (MSWD = 1.4) for Fish Canyon Tuff (Table 4).
Fish Canyon Tuff biotite produced a well-defined
plateau comprising thirteen steps forming 100% of the
released gas. Biotite from the Pagosa Peak Dacite
degassed much more abruptly, mainly in two steps. The
first seven steps and step 14, comprising only
approximately 4% of the total 39Ar released, were
excluded from the plateau (these steps have errors from
3.6 to 93 Ma, and removing them minimizes MSWD).
Both inverse isochrons gave results concordant with the
plateau ages (PPDlc = 28.19 ± 0.16 Ma, MSWD = 0.4;
FCTar = 28.19 ± 0.09 Ma, MSWD = 1.5) and atmospheric 36Ar/40Ar intercepts (Table 5 and Fig. 3).
Step-heating experiments on plagioclase also yielded
consistent ages of 28.26 ± 0.10 Ma (MSWD = 1.0) for
both the plateau age and the inverse isochron, with an
atmospheric 36Ar/40Ar intercept (Tables 4 and 5). The
highest temperature step of this analysis, however,
Table 4
Summary of step-heating 40Ar/39Ar age spectrum results on samples from the Fish Canyon magmatic system
Sample #
Material
K2O (wt.%)
K/Ca
a
39
Age Spectrum
NCD #1
NCD #2
MegaX
FCTar
FCTar
FCTar
FCTar
PPDcc
PPDlc
Sanidine
Sanidine
Sanidine
Sanidine
Biotite
Plagioclase
Hornblende
Sanidine
Biotite
11
11
11
11
9
1
1
11
9
72.2
70.1
41.2
70.4
26.1
0.2
0.1
64.6
21.5
7 of 11
10 of 11
8 of 24
12 of 12
13 of 13
11 of 12
9 of 11
12 of 13
6 of 14
99.7
99.9
70.1
100
100
93.5
97.8
99.1
96.3
28.05 ± 0.09
28.08 ± 0.09
28.23 ± 0.09
28.04 ± 0.09
28.19 ± 0.09
28.26 ± 0.10
27.49 ± 0.11
27.94 ± 0.09
28.22 ± 0.11
N
Ar (%)
MSWD
Age (Ma) ± 1σ
a
N = number of plateau steps used to calculate plateau age. See legend in Table 2 for details on analytical procedures.
0.5
2.6
0.3
2.0
1.4
0.9
6.0
2.3
0.7
144
O. Bachmann et al. / Chemical Geology 236 (2007) 134–166
Table 5
Summary of step-heating 40Ar/39Ar inverse isochron results on
samples from the Fish Canyon magmatic system
Sample # Material
40
Ar/36Ar ± 1σ Isochron analysis MSWD
Age (Ma) ± 1σ
NCD #1
NCD #2
MegaX
FCTar
FCTar
FCTar
FCTar
PPDcc
PPDlc
Sanidine
Sanidine
Sanidine
Sanidine
Biotite
Plagioclase
Hornblende
Sanidine
Biotite
529.5 ± 88.6
249.6 ± 83.9
294.0 ± 2.1
344.7 ± 18.3
296.3 ± 3.5
294.1 ± 15.3
297.1 ± 2.0
272.1 ± 57.2
291.0 ± 15.4
28.02 ± 0.09
28.06 ± 0.09
28.27 ± 0.09
28.01 ± 0.09
28.19 ± 0.09
28.26 ± 0.10
27.65 ± 0.14
27.96 ± 0.09
28.19 ± 0.16
0.8
0.5
0.8
1.1
1.5
1.0
1.1
2.4
0.4
The same number of steps was used to calculate both the plateau and
inverse isochron ages. See legend in Table 2 for details on analytical
procedures.
produced an age of more than 30 Ma and was excluded
from plateau age calculation. Thin-section observations
and microprobe analyses show that Fish Canyon
plagioclases contain cores with high anorthite contents
(up to An80), considerably exceeding the values
associated with clearly phenocrystic plagioclase (Bachmann et al., 2002). These calcic cores suggest recycling
from an earlier, more mafic stage of differentiation, and
may be responsible for this older step (see Layer and
Gardner (2001) for a similar interpretation), although
this has not been verified.
Hornblende yielded more complex results. Although
the three total-fusion ages (28.17 ± 0.16 Ma, 28.30 ±
0.16 Ma and 28.22 ± 0.14 Ma; average = 28.23 ±
0.12 Ma, MSWD = 0.2, Table 3) are identical to biotite and plagioclase ages at the 1σ level, the incremental heating experiment gave a younger plateau age of
27.49 ± 0.11 Ma with a large MSWD of 6.0, which
translates to 27.65 ± 0.14 Ma by the inverse isochron
method (MSWD = 1.1; Tables 4 and 5). The reason for
this discrepancy is not clear but, in light of the fact that
(1) the hornblende incremental heating age is younger
than all the other ages obtained on Fish Canyon minerals
and (2) hornblende release spectra during step-wise
heating in vacuo have proved to be complex due to the
structural decomposition of the heated grains (Lee et al.,
1991), we favor the total-fusion age and discard the
“anomalously” low incremental heating age.
3.3.3. Age probability analysis
The individual total-fusion results and all single-step
age data from incremental-heating experiments included
in the plateau calculations are shown in Figs. 4a
(sanidine from the three lithological units) and 5b
(sanidine, biotite, plagioclase and hornblende from the
Fish Canyon Tuff) in the form of age-probability
diagrams (Deino and Potts, 1992). The curves in these
diagrams sum the Gaussian frequency function values of
individual analyses, calculated from their means and
variances, at a range of age (abscissa) values. These
plots combine histogram-type distributions and analytical error information, and are ideally suited for
compilation and comparison of data series.
The probability curves for the sanidine data from
different samples overlap (Fig. 4a), but peak probability
for Pagosa Peak Dacite sanidine is shifted towards a
slightly younger age. Biotite does not seem to reproduce
this behavior, as Pagosa Peak Dacite biotite yields ages
that are very similar to those obtained for Fish Canyon
Tuff biotite on the basis of total fusion analyses
(PPDlc = 28.25 ± 0.10 Ma; FCTar = 28.25 ± 0.09 Ma),
weighted mean plateaus (PPDlc = 28.22 ± 0.11 Ma;
FCTar = 28.19 ± 0.09 Ma), and inverse isochrons
(PPDlc = 28.19 ± 0.16 Ma; FCTar = 28.19 ± 0.09 Ma).
The Fish Canyon Tuff sanidine peak is slightly younger
than the biotite and plagioclase peaks (Fig. 4b). The
smaller number of hornblende analyses produces a
poorly-defined peak that is nominally coincident with
the plagioclase-biotite group.
3.3.4. Feldspar megacryst
Feldspar megacrysts, up to several centimeters in
diameter, are present in pumices of the intracaldera Fish
Canyon Tuff (Fig. 5). These megacrysts, described in
Lipman et al. (1997) and Bachmann et al. (2002), record
a complex petrologic history. The cores are usually
relatively homogeneous K-feldspar, but the rims are
composite, showing two types of overgrowth textures:
(1) plagioclase mantling (rapakivi textures), developed
in response to changing conditions (most likely
temperature fluctuations) in the magma chamber
(Bachmann et al., 2002), and (2) granophyric rims
(fine-scale intergrowths of quartz and K-feldspar),
which apparently grew in response to a pressure drop
in the magma chamber associated with early eruptions
of the Fish Canyon magmatic system (Lipman et al.,
1997).
A step-heating analysis (24 steps) on one of these
megacrysts is used to investigate the potential presence
of argon memory in these complex crystals (Fig. 6). The
degassing pattern of this megacryst is divisible into three
groups of eight steps each. The first group of eight steps,
acquired at low laser power, shows a gradual rise from
approximately 27 Ma to more than 28 Ma. Their low
and variable K/Ca ratios indicate that these steps, which
contribute only a small amount to the total released 39Ar
(b9%), represent argon that was in part extracted from
O. Bachmann et al. / Chemical Geology 236 (2007) 134–166
145
Fig. 2. Incremental step-heating analyses of sanidine from the three units of the Fish Canyon magmatic system. Rectangle heights and all errors are
±1σ. Ages are relative to the Taylor Creek sanidine (TCs at 28.34 Ma; Renne et al., 1998).
plagioclase. The next eight steps produced the bulk of
the argon release and define a plateau (at 28.23 ±
0.09 Ma, MSWD = 0.3), which is slightly older than the
Fish Canyon sanidine age. This compares well, perhaps
coincidently, with the ages determined for biotite,
hornblende, and plagioclase. The last eight steps,
forming a significant portion of the gas release
(N20%), show a gradual rise from the plateau to almost
29 Ma. Their K/Ca ratios, which are very similar to
those of typical sanidine K/Ca ratios, may be interpreted
as gas release from domains in the K-feldspar core of the
megacryst. The 36Ar/40Ar intercept defined by the
inverse isochron of these last eight steps is atmospheric
within error, indicating that the 40Ar is a mix of closed-
system 40K decay and atmospheric sources (no evidence
for excess Ar). The age given by the 40Ar/39Ar intercept
on this inverse isochron is 28.75 ± 0.32 Ma (MSWD = 1.5;
Fig. 7).
3.4. Discussion of
40
Ar/39Ar data
High precision 40Ar/39Ar geochronology does not
discriminate between the eruption ages of the three units
of the Fish Canyon magmatic system. The time interval
over which they erupted was shorter than 0.3 Ma
(maximum time interval calculated by taking the oldest
and youngest mean sanidine ages of the stratigraphically
oldest and youngest sample (PPDcc and NCD #1,
146
O. Bachmann et al. / Chemical Geology 236 (2007) 134–166
Fig. 3. Incremental step-heating analyses of biotite, plagioclase and hornblende from the Fish Canyon Tuff and step-heating analysis from a Pagosa
Peak Dacite biotite. Errors and age calculations as for Fig. 2.
respectively) ± their 2σ errors; i.e., PPDcc = 27.96 +
0.18 Ma and NCD = 28.02–0.18 Ma). This estimate is
very conservative, and the interval was most likely
much shorter. Based on the absence of erosion or
deposition at the contact between PPD and FCT and the
lack of fumarolic alteration at the base of the Fish
Canyon Tuff, a more plausible time gap between the
Pagosa Peak Dacite and Fish Canyon Tuff would be on
the order of days to years (e.g., Riehle et al., 1995).
Our 40Ar/39Ar results on Fish Canyon sanidine
(inverse variance-weighted mean age of 28.02 ±
0.02 Ma (2σ; 102 out of 109 analyses; MSWD = 2.0)
calibrated to the an age of 28.34 Ma for Taylor Creek
Rhyolite; Renne et al., 1998) are in good agreement with
some published ages, wherein Fish Canyon sanidine
ages are calibrated against absolute timescales (i.e, U/Pb
ages, astronomical timescale, historical records): (1) the
intercalibration of U–Th–Pb and 40Ar/39Ar ages of
Villeneuve et al. (2000) gave an age of 27.98 ± 0.15 Ma
for the Fish Canyon sanidine, (2) an 40 Ar/ 39 Ar
experiment on Fish Canyon sanidine relative to sanidine
from the 79AD eruption of Vesuvius used as a fluence
monitor (Renne and Min, 1998) yielded 28.04 ±
0.45 Ma, and (3) the intercalibration of seven published
40
Ar/39Ar ages of polarity transitions measured relative
to Fish Canyon sanidine with the astronomically
O. Bachmann et al. / Chemical Geology 236 (2007) 134–166
147
Fig. 4. Bottom: age probability diagrams (Deino and Potts, 1992) of (a) all sanidine data from total fusion analyses and steps from incremental heating
experiments included in plateau calculations, and (b) of the four different mineral phases of the Fish Canyon Tuff. The ages reported correspond to the
maxima of the cumulative probability curves (input data at 1σ, excluding errors on J ). Top: ranked distributions of all single analyses included in the
calculations (error bars are 1σ and include the error on J ).
calibrated geomagnetic polarity time scale (Renne et al.,
1994) resulted in an age of 28.03 ± 0.18 Ma. It should be
mentioned, however, that another more recent astronomical calibration for the Fish Canyon sanidine (based
on the Cretan A1 ash layer) gave an age of 28.21 ±
0.04 Ma (1σ error, Kuiper et al., 2004).
Although analytical uncertainties preclude rigorous
conclusions, the age distribution obtained for the
different mineral phases (sanidine, plagioclase, biotite,
and hornblende) from the same sample (FCTar), and for
the feldspar megacryst, suggests the presence of
extraneous argon in the Fish Canyon minerals. If one
applies the widely accepted rule in geochronology that
two ages are different only if they do not overlap at 2σ
error, then all the 40Ar/39Ar ages obtained from the
different mineral phases are indistinguishable. The
incremental-heating experiment on the feldspar megacryst provides a more robust case for the presence of
extraneous argon, showing a 39Ar release spectrum
characterized by progressively older steps. A similar
Fish Canyon Tuff sanidine step-heating release spectrum (slight increase of age with increasing temperature)
has also been reported by Spell and McDougall (2003).
Two possibilities are commonly invoked to account
for seemingly old ages obtained by the 40Ar/39Ar
method: (1) either 40Ar is elevated due to excess argon
(the 40Ar/36Ar ratio of Ar added from external sources is
higher than the atmospheric value of 295.5, leading to
an excess in 40Ar) or (2) due to inherited radiogenic
argon (the dated material contains an component older
than the age of eruption). In the case of the Fish Canyon
system, several arguments suggest that the slightly older
ages are due to the presence of inherited argon in the
most retentive parts of the system.
1. Except for two sanidine analyses (NCD#1 and
FCTar), all inverse isochrons determined in this
study (even from the high temperature steps of the
megacryst incremental heating experiment) yield an
Fig. 5. Photomicrograph of a feldspar megacryst in a Fish Canyon Tuff
pumice from the intracaldera facies. The dashed line traces the
boundary between pumice and tuff matrix.
148
O. Bachmann et al. / Chemical Geology 236 (2007) 134–166
Fig. 6. Release spectrum and K/Ca ratio of a step-heating experiment on a Fish Canyon feldspar megacryst, compared with the step-heating pattern
from a Fish Canyon Tuff sanidine. The mean U–Pb age on Fish Canyon zircon (Oberli et al., 1990; Schmitz and Bowring, 2001) is also reported for
comparison. Errors and age calculations as in Fig. 2.
atmospheric 40Ar/36Ar intercept (295.5) within 2σ
errors.
2. “Staircase” release spectra, such as that in the
megacryst incremental-heating experiment, are typically produced by partially reset xenolithic material
(Gillespie et al., 1982, 1984; Heizler et al., 1999), and
are commonly interpreted as evidence for older argon
residing in the most retentive crystal lattice sites (e.g.,
Singer et al., 1998). Although the ages of individual
steps may not be distinguishable from one another,
the four incremental-heating experiments on sanidine
reported in Fig. 2 (and the step-heating age spectrum
of Spell and McDougall, 2003; their Fig. 2) show
similar “staircase” release spectra, adding some
weight to the idea of inherited argon in the most
retentive lattice sites.
3. Excess argon tends to be relatively uncommon in
minerals from silicic volcanic rocks (for an exception
to this, see Layer and Gardner, 2001), largely because
argon is highly incompatible in all major igneous
minerals (Kelley, 2002). However, excess Ar can
occur when a significant volume fraction of melt
(and/or fluid inclusions) is present in the analyzed
minerals (inclusions can contain up to 100–1000
times more argon than their host; Kelley, 2002). This
inclusion-derived excess 40Ar could be significant in
the Fish Canyon magma as nearly all mineral phases
are known to contain some melt inclusions.
Fig. 7. Inverse isochrons obtained from the step-heating experiment on the Fish Canyon feldspar megacryst, including (a) all 24 steps, and (b) only the
last eight steps of the analysis. Both have atmospheric 36Ar/40Ar intercepts. Errors and age calculations as in Fig. 2.
O. Bachmann et al. / Chemical Geology 236 (2007) 134–166
However, if this were the case, one might expect the
sanidine crystals to give the oldest 40Ar/39Ar ages
(sanidine contains the most abundant and largest melt
inclusions of all analyzed crystalline phases; Bachmann et al., 2002), although the elevated K content of
sanidine (as compared to low-K minerals such as
plagioclase and hornblende) might partially mask this
effect. The absence of a correlation between older ages
and the presence of melt inclusion weakens the case for
excess 40Ar being derived from such a reservoir.
Evidence for old (inherited?) argon in plagioclase,
hornblende and biotite has also been reported for the
∼ 2800 km3 Young Toba Tuff (YTT), an ignimbrite with
many petrologic similarities to the Fish Canyon Tuff
(Chesner, 1998; Gardner et al., 2002; Thomas et al.,
2003). Relative to the narrow peak defined by sanidine
at 74 ± 4 ka (the eruption age), biotite, hornblende, and
to a lesser extent plagioclase, show skewed 40Ar/39Ar
age distributions, which predate eruption by up to
1.5 Ma (Gardner et al., 2002; Thomas et al., 2003).
The apparent presence of inherited argon in minerals
of the Fish Canyon Tuff (and Young Toba Tuff) leads to
questions concerning argon diffusion in silicate minerals. The mineral textures and zoning patterns in the Fish
Canyon Tuff, as well as water and heat budgets, have
been interpreted in terms of a protracted crystallization
history and extended residence (possibly N105 years) at
supra-solidus temperatures (≥ 700 °C; Bachmann and
Dungan, 2002; Bachmann et al., 2002; Bachmann and
Bergantz, 2003). On the other hand, closure temperatures calculated for volume diffusion (Dodson, 1973)
predict that, under these conditions, every major mineral
phase in these magmas should have remained fully open
to Ar loss prior to eruption. To explain the presence of
inherited argon in magmas, Gansecki et al. (1996),
Singer et al. (1998), and Gardner et al. (2002) have
suggested that the incompletely reset minerals were
xenocrysts with short residence times (10–100 years).
This idea is attractive for relatively small, hot magma
bodies (Singer et al., 1998), but appears doubtful in the
case of voluminous, crystal-rich magmas at near-solidus
temperatures, such as those tapped by the Toba and Fish
Canyon eruptions. A truly xenocrystic origin would
imply large amounts of assimilation of upper crust
followed by unrealistically rapid dissemination of solid
material in chambers filled with high-viscosity magmas,
which are unlikely to convect turbulently. Alternatively,
argon may be retained in minerals at higher temperatures or for much longer periods than those predicted by
volume diffusion and currently available diffusion rates
estimates. Diffusion in minerals is arguably a complex
149
phenomenon, even in gem-quality crystals (Wartho
et al., 1999), and activation energies may differ
substantially for different crystal domains, allowing
trapping of argon in “retentive” parts of crystals even at
magmatic temperature (Foland, 1994).
4. Zircon U–Pb dating
4.1. Sample description
Samples for U–Pb dating were selected from
multiple localities on the basis of our refined eruptive
stratigraphy for the Fish Canyon system (Lipman et al.,
1997; Bachmann et al., 2000) and for correspondence
with the 40Ar/39Ar study. Three of the samples that were
investigated in our 40Ar/39Ar study were also dated by
U/Pb (FCTar — basal vitrophyre of outflow Fish
Canyon Tuff, PPDcc — glassy Pagosa Peak Dacite,
and NCD — Nutras Creak Dacite). The FCT outflow
sample (FCTfv) collected by the USGS to serve as an
inter-laboratory standard for the 40Ar/39Ar and fissiontrack techniques (Neaser et al., 1981) and studied by
Oberli et al. (1990) is also reported.
In addition, four holocrystalline xenoliths from two
different localities in the intracaldera FCT (see Fig. 1)
were also dated. GrnX is a granitic xenolith (∼76 wt.%
SiO2), with a mineral assemblage dominated by quartz
and K-feldspar (N70% of the rock). Plagioclase, biotite
and Fe–Ti oxides form the remaining 30%, in roughly
equal proportions (∼10% each). The texture is bimodal
(porphyritic), with millimeter-sized grains of plagioclase
and biotite in a finer matrix (∼100 μm) of quartz, Kfeldspar, plagioclase, biotite and oxides. GrdX1 and
GrdX2 are granodioritic xenoliths (∼68 wt.% SiO2), with
mineral assemblages identical to the Fish Canyon magma
(Pl + Kfs + Qtz + Hbl + Bt + Spn + Mag + Ilm + Ap + Zrn).
The modal abundances of hydrous minerals (hornblende
and biotite), Fe–Ti oxides and titanite are also comparable
to the Fish Canyon magma, but GrdX1 and GrdX2
contain higher proportions of quartz, plagioclase, and Kfeldspar, reflecting the absence of glass in these
holocrystalline samples. The major mineral phases are
equigranular and coarse-grained (0.5–5 mm), with
generally euhedral hornblende, biotite, and plagioclase,
but anhedral (interstitial) Kfs and Qtz. TonX is a tonalitic
xenolith that is slightly less silicic than GrdX1 and GrdX2
(∼66 wt.% SiO2) but is texturally and mineralogically
similar, except for a higher modal proportion of
plagioclase. All these xenoliths are fresh and show no
signs of deformation.
The zircon populations of these eight rocks have many
features in common. The colorless to pale amber-colored
150
O. Bachmann et al. / Chemical Geology 236 (2007) 134–166
grains are prismatic (when euhedral), with aspect ratios of
∼1.5–6. The external morphology of the crystals centers
around S17–S18 based on the typological classification
scheme of Pupin (1980); i.e., the (100) prism and (211)
pyramidal faces slightly dominate the (110) and (101)
faces. Fine to coarse mineral and glass (melt) inclusions
are present in most of the grains selected for U–Pb
analysis (with concomitant effects on common lead
content, see below). The melt inclusions are of variable
geometry, ranging from blobs to vermicular and elongate
shapes. Some of the elongate melt inclusions, when
oriented parallel to the c axis, are reminiscent of initial
skeletal zircon growth under supersaturation conditions.
Due to the presence of intergrown phases and abundant
negative mineral imprints, the external faces and edges of
many zircons have a rather rough and often irregular
(anhedral) appearance. The quality of the external faces
of the zircons is thus highly variable, not only within, but
also between the analyzed populations. Whereas the preFCT Pagosa Peak Dacite (PPDcc) and the post-FCT
Nutras Creek Dacite (NCD) show a relatively large
number of euhedral grains, the population of FCTar (FCT
outflow facies) is dominated by grains with irregular
surfaces. The remaining samples yielded zircons with
intermediate morphological characteristics. FCTar is
comparatively poor in zircon and its population is
characterized by small average grain width (b 80 μm),
whereas NCD (particularly rich in zircon), PPDcc, GrnX,
GrdX1, GrdX2, and TonX contain a higher proportion of
grains N120 μm.
Fig. 1 of Schmitz and Bowring (2001) shows typical
aspects of growth textures for Fish Canyon Tuff zircons.
Cathodoluminescence images available for zircons from
samples FCTar, PPDcc, GrnX, GrdX1 and GrdX2 reveal
similar textures. All crystals display thick mantles of
concentric, fine oscillatory zoning. Whereas in some of
the zircons this regular zoning pattern is present from
core to rim, the inner parts of the majority of the crystals
are characterized by broader zones and domains, which
often display highly irregular, partly embayed boundaries and can occasionally be observed to cut concentric
zonation. Melt inclusions are typically surrounded by
prominent, brightly luminescent coronas forming
embayments in the oscillatory-zoned parts of the
crystals. Coronas surrounding melt inclusions located
in the centers of crystals have in some cases developed
into bright coherent core domains exhibiting euhedral
boundaries. Inclusions located at more eccentric positions often result in growth impedance and re-entrant
crystal faces propagating to the surface of the grain.
Most zircons display one or more dark, narrow, euhedral
zones that occur at variable distances from crystal cores
(e.g., Fig. 1a and d of Schmitz and Bowring, 2001).
Corrosion features at intermediate positions within the
crystals are mainly confined to rounding of pyramidal
terminations. Based on isotopic evidence, a small
number of older inherited cores are likely to be present
in the imaged populations, but cannot be recognized
unequivocally due to the absence of typical morphological indicators and zonal growth distortion induced
by abundant inclusions.
4.2. Analytical procedures
Depending on sample size, individual rock specimens weighing from 0.08 to 5.4 kg were crushed by
hydraulic press, jaw crusher or hammer and reduced
tob315 μm particle size by use of a disk mill or a
swinging-disk mill. Heavy-mineral concentrates were
obtained using a Wilfley table in the case the largest
sample (GrnX) and/or standard methyl-iodide heavyliquid separation techniques. If required, zircon was
further concentrated using a Frantz Isodynamic Magnetic Separator or Clerici's solution.
4.2.1. Methods used at ETH Zurich
Analyses 1–37 (Table 6) were performed at ETH
Zurich. In order to minimize potential bias created by
the presence of restitic cores, pre-selected zircon grains
were mounted in glycerol and studied by transmitted
light microscopy for detection of inherited components, which are often revealed by minute fluid inclusions (“bubble” cores) and other mineral impurities
at the core/host interface. Except for the zircons from
sample FCTfv (zircons 9–15, Table 6), all crystals
have been air abraded (Krogh, 1982) in an Al2O3 dish
for removal of surface contaminants, potentially altered parts, and mineralogical impurities attached to
the grains.
U–Pb analysis of zircons followed methods described by Meier and Oberli in Wiedenbeck et al.
(1995). All analyses were spiked using the same mixed
233
U–235U–230Th–205Pb tracer. U/Pb in this tracer has
been calibrated to better than 0.1% as confirmed by
repeat calibrations ab inito and by an independent
recalibration in another laboratory (Mundil et al., 2001).
Because of small sample weights and enhanced
solution/solute ratios, a re-equilibration step using HCl
following bomb decomposition was omitted. Lead and
uranium isotopic measurements on samples 9–15 were
performed on a Varian MAT Tandem mass spectrometer
using a highly linear secondary-electron multiplier in
low-gain analog mode, correcting multiplier-related
mass discrimination by multiplication of the isotopic
O. Bachmann et al. / Chemical Geology 236 (2007) 134–166
ratios with the square-root of the corresponding mass
ratios in addition to corrections related to mass
fractionation by the thermal ionization/evaporation
process. All the remaining samples were measured on
a Finnigan MAT 262 RPQplus multiple-collector mass
spectrometer using single ETP AF 150H or ETP DM198
multipliers in pulse-counting mode. Whereas earlier
generations of AF 150H multipliers have shown good
linearity requiring minimal correction (see Oberli et al.
(1999) for details), more elaborate corrections have
become necessary for later series. These corrections are
based on repeat calibration runs using the 207 Pb/206Pb
and 204Pb/206Pb ratios (normalized by 208Pb/206Pb)
from the NIST SRM 982 equal atom standard at count
rates of 206Pb varying from ∼ 0.1 to 1 MHz. Contribution from dark noise to the ion signals has been
monitored during all runs and has been found to be
negligible (the dark count rates were typically
≪ 1 cpm). Whereas total procedural blanks for Pb,
based on multiple blank experiments, were relatively
high for an early (1987 to 1990) series of analyses, the
blanks have since stabilized at 0.9–1.5 pg for all more
recent analyses (for details see caption of Table 6).
All analytical uncertainties quoted in Table 6, the text
and in figures, unless stated otherwise, are given at
the 95% confidence level. Error estimation is based on
the Gaussian error propagation scheme documented
in the Appendix. It includes the uncertainties of mass
spectrometric ratio measurements, reproducibility of
instrumental mass fractionation (Pb only; uncertainties
assigned to the measured U contents are calculated
directly from the double-spike procedure) and reproducibility of Pb amounts and isotopic composition measured
in several series of total blank experiments (these include
covariance terms between amounts of Pb in the blanks
and isotopic ratios). No corrections have been made for
U blanks, which were negligible. In order to also cover
the less than ideal behavior of the current ion counting
system, we have repeatedly analyzed variable amounts
of a mixed U–Pb solution containing 5–20 pg of radiogenic Pb prepared from NIST SRM 960 and SRM
983 reference materials and have determined an
external variance component (Oberli et al., 1999) of
0.12% (1σ) for 206 Pb/ 238 U in excess of expected (propagated) analytical errors. This excess variance component has been propagated into the uncertainty of the
206
Pb/ 238 U ratio in order to avoid underestimation of
the error. The uncertainty in the isotopic composition of
common lead inherited by these samples has not been
included in the error propagation procedure as we
consider this component to be a systematic rather than a
stochastic variable. Isotopic ratios used for common
151
lead corrections are listed in Table 6 and discussed in
Section 4.4.2.
4.2.2. Methods used at the Berkeley Geochronological
Center (BGC)
The substantial spread of the U–Pb ages in excess of
analytical error observed for the zircon samples
measured at ETH Zurich (see Section 4.3.1 below)
prompted us to analyze complementary individual
crystals using the pretreatment of thermal annealing
combined with chemical abrasion (“CA-TIMS” method;
Mattinson, 2005, Mundil et al., 2004). Selected zircons
from the Pagosa Peak Dacite (PPD) were first annealed
at 850 °C for 48 h and then etched in a 29 M HF solution
at 220 °C in pressurized dissolution capsules for 16 h.
U–Pb analyses were then performed on the etched
residues of five single grains. For comparison, additional five analyses were carried out on single zircons
without any pretreatment. For a more detailed description of analytical techniques and data reduction
procedures used at BGC the reader is referred to Mundil
et al. (2004) and Table 7.
4.3. Results
4.3.1. ETH results
The results of 37 U–Pb analyses are listed in Table 6
and displayed in Figs. 8 and 9. Analyses 9–15 (FCTfv,
UGS fission-track standard sample) were performed
during 1987–1990 (Fischer et al., 1989; Oberli et al.,
1990), whereas the remaining data have been determined more recently. Except for three of these early
determinations (9–11), for which 6 to 7 crystals were
analyzed, the data represent measurements on single
crystals. The 206 Pb/238 U ratios and ages displayed in
Fig. 8 and Table 6 have been corrected for initial
radioactive disequilibrium in 230Th / 238 U (Schärer,
1984) by adopting Th/U = 2.2 (Schmitz and Bowring,
2001) for the host magma (see discussion in Section
4.4.2).
When viewed in 206Pb/238U versus 207Pb/235U space
(Fig. 8) all results except those for zircon PPDcc
8 (Fig. 8h) are analytically concordant. The 206Pb/238 U
ages of 32 out of the remaining 36 analyses fall within
the time interval of 28.0 to 28.7 Ma (main-group dates).
Three zircons from a tonalitic xenolith in intracaldera
FCT (TonX 24–26; Fig. 8j) show reproducible ages
with a weighted mean at 31.31 ± 0.05 Ma (95% c.l.
internal; MSWD = 0.8), predating the main-group
dates. Furthermore, one of four analyses performed on
zircons from the porphyritic granitic xenolith collected
from the FCT (GrnX 27) also gave an older age (30.29 ±
152
O. Bachmann et al. / Chemical Geology 236 (2007) 134–166
Table 6
Summary of U–Pb results on zircons from the Fish Canyon magmatic system measured at ETH
Zircona
I.D
Weightb
[μg]
U
[ppm]
Th/Uc
[wt]
Pbradd
[ppm]
Pbcome
206
208
207
206
204
206
235
238
Pb/
Pbf
Pb/
Pbg,h
Pb/
Ug g
Pb/
U g,h
ρi
[pg]
207
206
235
238
Age [Ma]
Age [Ma]
Pb/
U
Pb/
Uh
PPDcc: Pagosa Peak Dacite (pre-FCT pyroclastic unit)
1
12.8
446
0.56
2.1
2.8
562
2
8.8
164
0.83
0.8
1.7
251
3
9.8
354
0.51
1.6
1.0
945
4
10.4
382
0.56
1.8
2.8
394
5
4.6
322
0.58
1.5
2.9
157
6
4.5
147
1.24
0.8
1.2
160
7
3.5
376
0.79
1.9
1.1
440
8
3.7
377
0.68
2.9
2.0
317
0.1833 ± 7
0.2748 ± 18
0.1694 ± 7
0.1856 ± 9
0.1900 ± 23
0.4084 ± 30
0.2610 ± 25
0.2233 ± 14
0.02840 ± 16
0.02841 ± 51
0.02847 ± 19
0.02824 ± 25
0.02833 ± 61
0.02809 ± 81
0.02817 ± 90
0.04859 ± 52
0.004413 ± 12
0.004421 ± 34
0.004426 ± 19
0.004408 ± 14
0.004398 ± 13
0.004358 ± 14
0.004429 ± 15
0.007013 ± 20
0.30
0.48
0.51
0.32
0.38
0.47
0.26
0.26
Mean Age j:
MSWD:
28.44 ± 16
28.44 ± 51
28.51 ± 19
28.28 ± 25
28.36 ± 60
28.13 ± 80
28.21 ± 89
48.17 ± 51
28.42 ± 10
0.5
28.39 ± 8
28.44 ± 22
28.47 ± 12
28.35 ± 9
28.26 ± 9
28.03 ± 9
28.49 ± 9
45.05 ± 13
28.33 ± 15
11.4
FCTfv: Fish Canyon Tuff (outflow facies, Fun Valley location)
9 (7)
165.5
314
0.59
1.5
51.3
298
10 (6)
154.6
328
0.60
1.6
49.6
300
11 (6)
91.5
352
0.54
1.6
41.9
230
12
34.1
397
0.49
1.8
15.7
257
13
22.5
209
0.61
1.0
11.4
131
14
21.8
372
0.49
1.7
8.9
267
15
9.1
403
0.57
1.9
6.8
163
0.1955 ± 13
0.1978 ± 22
0.1762 ± 25
0.1625 ± 48
0.2021 ± 14
0.1607 ± 31
0.1861 ± 51
0.02844 ± 28
0.02825 ± 62
0.02886 ± 64
0.0287 ± 12
0.0284 ± 35
0.02761 ± 84
0.0280 ± 14
0.004417 ± 15
0.004413 ± 19
0.004415 ± 20
0.004457 ± 20
0.004431 ± 37
0.004406 ± 16
0.004404 ± 18
0.38
0.41
0.42
0.58
0.85
0.47
0.61
Mean Age j:
MSWD:
28.47 ± 28
28.29 ± 61
28.89 ± 63
28.7 ± 12
28.4 ± 35
27.66 ± 83
28.0 ± 14
28.44 ± 22
1.1
28.41 ± 10
28.38 ± 13
28.40 ± 13
28.67 ± 13
28.50 ± 24
28.34 ± 10
28.33 ± 12
28.41 ± 10
3.4
FCTar: Fish Canyon Tuff (outflow facies, Agua Ramon location)
16
9.4
446
0.55
2.1
10.0
134
0.1795 ± 16
17
6.6
293
0.54
1.4
1.8
304
0.1766 ± 13
18
12.2
459
0.51
2.1
4.3
372
0.1671 ± 7
19
17.9
227
0.78
1.1
7.8
163
0.2556 ± 17
0.02830 ± 36
0.02885 ± 36
0.02842 ± 17
0.02856 ± 39
0.004400 ± 17
0.004449 ± 14
0.004427 ± 12
0.004431 ± 13
0.41
0.32
0.31
0.38
Mean Age j:
MSWD:
28.34 ± 36
28.88 ± 35
28.45 ± 17
28.59 ± 39
28.51 ± 13
2.0
28.30 ± 11
28.62 ± 9
28.47 ± 8
28.50 ± 8
28.49 ± 18
6.9
NCD: Nutras Creek Dacite (post-FCT lava flow)
20
5.6
319
0.86
1.6
1.8
21
7.4
218
0.74
1.1
3.0
22
12.4
252
0.83
1.2
4.2
23
9.9
228
0.78
1.1
1.8
286
162
221
353
0.2837 ± 19
0.2425 ± 25
0.2743 ± 16
0.2560 ± 15
0.02863 ± 56
0.02813 ± 71
0.02803 ± 40
0.02841 ± 44
0.004425 ± 14
0.004389 ± 14
0.004372 ± 12
0.004434 ± 15
0.29
0.39
0.32
0.28
Mean Age j:
MSWD:
28.66 ± 55
28.16 ± 70
28.07 ± 40
28.44 ± 43
28.31 ± 24
1.2
28.46 ± 9
28.23 ± 9
28.12 ± 8
28.52 ± 10
28.31 ± 30
18.2
TonX: tonalitic xenolith in intracaldera Fish Canyon Tuff
24
11.5
169
0.47
0.9
4.4
149
25
10.6
306
0.53
1.6
3.1
323
26
9.6
232
0.48
1.2
1.7
390
0.1560 ± 14
0.1735 ± 11
0.1571 ± 8
0.03117 ± 36
0.03163 ± 31
0.03136 ± 22
0.004862 ± 16
0.004877 ± 16
0.004868 ± 14
0.36
0.35
0.29
Mean Age j:
MSWD:
31.17 ± 36
31.62 ± 31
31.35 ± 22
31.39 ± 16
1.9
31.27 ± 10
31.36 ± 10
31.30 ± 9
31.31 ± 5
0.8
GrnX: Porphyritic granitic xenolith in intracaldera Fish Canyon Tuff
27
14.5
360
0.38
1.7
4.2
381
0.1248 ± 16
28
7.7
284
0.68
1.4
3.5
188
0.2244 ± 18
29
7.7
227
0.68
1.1
2.1
245
0.2223 ± 17
30
12.5
239
0.72
1.2
6.2
148
0.2375 ± 19
0.03035 ± 30
0.02830 ± 48
0.02862 ± 58
0.02814 ± 50
0.004709 ± 17
0.004418 ± 14
0.004435 ± 16
0.004381 ± 14
0.34
0.34
0.36
0.40
Mean Age j:
MSWD:
30.36 ± 29
28.34 ± 47
28.65 ± 58
28.18 ± 50
28.36 ± 29
0.8
30.29 ± 11
28.42 ± 9
28.53 ± 10
28.18 ± 9
28.37 ± 44
14.6
O. Bachmann et al. / Chemical Geology 236 (2007) 134–166
153
Table 6 (continued )
Zircona
I.D
Weightb
[μg]
Th/Uc
U
[ppm]
[wt]
Pbradd
[ppm]
Pbcome
206
208
207
206
204
206
235
238
Pb/
Pbf
Pb/
Pbg,h
Pb/
Ug g
Pb/
U g,h
ρi
[pg]
207
206
235
238
Age [Ma]
Age [Ma]
Pb/
U
Pb/
Uh
GrdX1: granodioritic xenolith in intracaldera Fish Canyon Tuff
31
2.4
931
0.95
4.8
3.4
192
0.3140 ± 21
32
7.6
601
0.58
2.8
6.7
206
0.1904 ± 14
33
4.5
307
0.64
1.5
3.7
119
0.2107 ± 30
0.02860 ± 58
0.02848 ± 32
0.02823 ± 84
0.004407 ± 13
0.004417 ± 12
0.004378 ± 14
0.36
0.32
0.43
Mean Agej:
MSWD:
28.63 ± 57
28.51 ± 32
28.27 ± 83
28.51 ± 26
0.3
28.34 ± 8
28.41 ± 8
28.16 ± 9
28.31 ± 31
8.9
GrdX2: Granodioritic xenolith in intracaldera Fish Canyon Tuff
34
1.3
294
0.79
1.5
1.5
87.8 0.2593 ± 74
35
1.5
569
0.62
2.7
3.6
83.3 0.2042 ± 42
36
1.4
486
0.55
2.3
2.3
96.7 0.1809 ± 46
37
0.7
269
0.62
1.3
8.0
24.9
0.204 ± 30
0.0272 ± 24
0.0279 ± 12
0.0286 ± 15
0.0288 ± 77
0.004403 ± 26
0.004426 ± 17
0.004399 ± 28
0.004402 ± 55
0.47
0.44
0.39
0.86
Mean Agej:
MSWD:
27.3 ± 24
28.0 ± 12
28.6 ± 15
28.8 ± 76
28.12 ± 85
0.3
28.32 ± 17
28.47 ± 11
28.30 ± 18
28.32 ± 35
28.40 ± 15
1.3
a
Single zircon analyses except numbers enclosed by parentheses correspond to the number of crystals analyzed in the case of multi-grain samples.
Weighed to a precision of ± 0.15 μg (1σ reproducibility) using an ultra-micro balance.
c
Calculated from radiogenic 208Pb/206Pb adopting t = 28.4 Ma for all samples except TonX, for which t = 31.3 Ma has been used.
d
Concentration of radiogenic lead in sample.
e
Content of common lead in analysis (includes analytical blank).
f
Measured ratio corrected for mass fractionation only.
g
Radiogenic lead. Analytical uncertainties are given at the 95% confidence level and refer to the least significant digits of the corresponding
values. Data corrected for mass fractionation, analytical blank and sample common lead using the parameters given below.
h
Data corrected for initial radioactive disequilibrium in 230Th /238U.
i
Correlation coefficient of 207Pb /235U versus 206Pb/238U.
j
Weighted mean 207Pb/235U ages and206Pb/238U ages, omitting analyses 8 and 27. Errors are 95% c.l. external for MSWD N 2.0 and 95% c.l.
internal for MSWD ≤ 2.0 (MSWD N 2.0 forces use of external errors for the given data sets, with probability-of-fit b 10%).
Parameters used for data reduction and error propagation (reproducibility given at 1σ level):
b
– Pb mass fractionation correction factor: 1.0009 ± 0.0004 amu− 1 (samples 9–15, measured 1987 and 1990) and 1.0005 ± 0.0004 amu− 1 (remaining
samples measured 2000 and later); U fractionation correction by double-spiking techniques.
– Blank data used for individual samples or group of samples (the numbers denoting reproducibility refer to the least significant digits of the
corresponding values):
I.D.
Pbtot [pg]
208
207
204
ρ (207Pb/206Pb vs. 204Pb/206Pb)
ρ (206Pb vs. 204Pb/206Pb)
1–8, 16–19, 27–30
9–11
12
13
14–15
20–23, 31–37
24–26
0.86 ± 0.11
17.7 ± 3.5
10.2 ± 1.9
10.8 ± 1.9
5.9 ± 1.8
0.93 ± 0.17
1.53 ± 0.10
2.0450 ± 34
2.0532 ± 98
2.076 ± 12
2.0888 ± 89
2.088 ± 10
2.0475 ± 85
2.0831 ± 39
0.8359 ± 18
0.8453 ± 55
0.8549 ± 69
0.8624 ± 68
0.8645 ± 70
0.8396 ± 67
0.8625 ± 24
0.05339 ± 21
0.05413 ± 63
0.0547 ± 12
0.0554 ± 13
0.05519 ± 60
0.05340 ± 24
0.05511 ± 17
−0.94
0.54
0.53
0.52
0.59
−0.61
0.95
− 0.15
∼0
− 0.28
− 0.06
− 0.09
− 0.09
0.70
Pb/206Pb
Pb/206Pb
Pb/206Pb
U blanks were negligible (no corrections required).
– Initial Pb composition: 206Pb/ 204Pb = 18.448, 207Pb/ 204Pb = 15.569, 208Pb / 204Pb = 37.665 (mean values representative for plagioclase from Fish
Canyon Tuff, Riciputi et al., 1995).
0.11 Ma). The three other dates from this xenolith
conform to those of the main group (Fig. 8i).
For ease of comparison, Fig. 8a–g show all maingroup dates obtained for the individual rock samples
using identical scaling. We note that six out of seven
main-group populations are characterized by ranges of
ages that are in excess of analytical error for the precise
206
Pb/238U ratios, as indicated by non-overlapping error
ellipses and by elevated MSWD values (Table 6). This is
also illustrated in Fig. 9a, which displays the ranked
206
Pb/238U ages of all main-group zircons including their
1σ error. These 206Pb/238U ages range from 28.67 ±
0.13 Ma (FCTfv 12) to 28.03 ± 0.09 Ma (PPDcc 6), the
two dates being different by 0.64 ± 0.16 Ma. The inverse-
154
O. Bachmann et al. / Chemical Geology 236 (2007) 134–166
Table 7
U–Pb results on zircons from the Pagosa Peak Dacite (PPD) measured at BGC
Zircona I.D.
Weightb U
Th/Uc Pbradd Pbcome
[μg]
[ppm]
[ppm] [pg]
BGC.Z01CA 10.4
BGC.Z02CA 8.9
BGC.Z03CA 4.4
BGC.Z04CA 1.3
BGC.Z05CA 2.0
BGC.Z08
2.8
BGC.Z09
5.7
BGC.Z10
3.0
BGC.Z11
2.2
BGC.Z12
3.2
140
310
141
260
299
624
299
884
961
1138
0.78
0.52
0.86
0.51
0.57
0.68
0.76
0.65
0.56
0.56
0.7
1.4
0.7
1.2
1.4
3.0
1.5
4.2
4.5
5.3
5.9
2.1
1.8
1.8
2.8
3.1
2.7
2.9
2.8
10.7
206
208
207
206
ρi
207
88
396
117
74
80
180
191
281
226
112
0.2483 ± 08
0.1645 ± 05
0.2726 ± 09
0.1614 ± 19
0.1818 ± 11
0.2169 ± 09
0.2422 ± 07
0.2070 ± 07
0.1780 ± 06
0.1780 ± 11
0.0288 ± 49
0.0286 ± 19
0.0291 ± 38
0.0290 ± 62
0.0315 ± 63
0.0291 ± 22
0.0276 ± 21
0.0285 ± 14
0.0285 ± 17
0.0289 ± 37
0.004429 ± 41
0.004445 ± 17
0.004453 ± 36
0.004462 ± 53
0.004491 ± 54
0.004426 ± 19
0.004381 ± 31
0.004438 ± 13
0.004404 ± 31
0.004418 ± 34
0.89
0.93
0.78
0.88
0.91
0.78
0.52
0.82
0.49
0.83
28.8 ± 4.9
28.6 ± 1.9
29.1 ± 3.8
29.0 ± 6.2
31.5 ± 6.3
29.1 ± 2.3
27.6 ± 2.1
28.5 ± 1.4
28.6 ± 1.8
29.0 ± 3.7
Pb/204Pbf
Pb/206Pbg,h
Pb/235Ug
Pb/238Ug,h
Pb/235U 206Pb/238U
Age [Ma] Age [Ma]
28.49 ± 27
28.59 ± 11
28.64 ± 23
28.70 ± 34
28.89 ± 35
28.47 ± 12
28.18 ± 20
28.55 ± 08
28.33 ± 20
28.42 ± 22
Uncertainties of individual ratios and ages are given at the 2σ level and do not include errors of decay constants.
Repeat measurements of the total procedural blank averaged 0.8 ± 0.3 pg Pb (U blanks were indistinguishable from zero), with 206Pb/204Pb = 18.55 ±
0.63, 207Pb/204Pb = 15.50 ± 0.55, 208Pb/204Pb= 38.07 ± 1.56 (all 2σ of population), and a 206Pb/204Pb–207Pb/204Pb correlation of +0.9.
a
CA denotes application of thermal annealing combined with chemical abrasion (Mundil et al., 2004; Mattinson, 2005).
b
sample weight is calculated from crystal dimensions and is associated with as much as 50% uncertainty (estimated).
c
present day Th/U ratio calculated from radiogenic 208Pb/206Pb and age.
d
concentration of radiogenic Pb in sample.
e
content of common Pb in analysis (includes analytical blank).
f
corrected for tracer contribution and mass fractionation (0.15 ± 0.09%/amu).
g
radiogenic Pb; data corrected for mass fractionation, tracer contribution and common Pb contribution (see below).
h
Data corrected for initial radioactive disequilibrium in 230Th/238U adopting Th/U = 2.2 for the crystallization environment.
i
correlation coefficient of radiogenic 207Pb/235U versus 206Pb/238U.
variance weighted mean 206 Pb/238 U age of the 32 analyses is 28.369 ± 0.052 Ma (95% c.l., external error).
MSWD = 8.4 corroborates the observation that the
cumulative range of ages significantly exceeds the
amount of variation that could be attributed to analytical
error. There is essentially zero probability that the ages
are distributed relative to a common mean. Because
MSWD = varext / varint, where varext and varint respectively denote external and internal variance of the mean,
an MSWD value of 8.4 indicates that 88% of the external
variance is due to a real spread in ages.
The mean weighted 207 Pb/235 U age is 28.439 ±
0.063 Ma (95% c.l., external error) and marginally
overlaps with the mean 206Pb/238U age. Because of very
low radiogenic 207 Pb contents, hence unfavorable
207
Pbrad/207 Pbcom ratios, the analytical errors of the
individual 207Pb/235 U ages are on average ∼ 12 times
larger than the corresponding 206Pb/238 U age errors
(disregarding analyses FCTfv 12, 13 and 15, which have
errors N2 Ma). In contrast to the latter, an excess
variance cannot be resolved from the distribution of the
207
Pb/235U ages due to dominance of the internal
(analytical) error component (this also holds for the
individual data sets, as can be seen from comparison of
the MSWD values listed in Table 6). MSWD = 0.21 even
suggests that the analytical errors assigned to the
207
Pb/235U dates may have been overestimated. Further
discussion will therefore focus on the more robust
206
Pb/238U dates.
4.3.2. BGC results
Because the variation in U–Pb ages obtained for the
Pagosa Peak Dacite (PPD) alone covered a large fraction
(28.49 ± 0.09 Ma to 28.03 ± 0.09 Ma, excluding zircon
8) of the total spread in ages, zircons from this sample
were chosen for further tests aimed at resolving two
questions. These are; 1) could the spread of the ages be
reproduced by an independent analyst, and, if this were
the case, 2) does the application of novel preparation
techniques (“chemical abrasion”, Mattinson, 2005;
Mundil et al., 2004) yield further information on the
systematics of the observed age distribution?
The results of 10 analyses on single grains and
fragments of single grains are listed in Table 7 and
shown in Fig. 10. The disequilibrium corrected (+ 65–
80 ka) 206Pb/238U ages for untreated zircons spread
from 28.55 ± 0.08 Ma to 28.18 ± 0.20 Ma and give a
weighted mean age of 28.46 ± 0.15 Ma. MSWD = 3.5
indicates scatter in excess of analytical error. A weighted
mean 207Pb/235U age of 28.50 ± 0.85 Ma is less precise
due to the higher impact of common Pb correction but
agrees within uncertainty. Both the distribution of the
individual ages and the mean age are in overlapping
relationship with those obtained at ETH.
O. Bachmann et al. / Chemical Geology 236 (2007) 134–166
155
Fig. 8. U–Pb concordia diagrams for all zircon populations analyzed in this study. Identical scaling has been used for all diagrams except for (h), (i)
and ( j), which include age results N29 Ma.
Samples pretreated by CA have higher U–Pb ages
than those analyzed without any pretreatment. The
pretreated zircons yield a coherent weighted mean age
of 28.61 ± 0.08 Ma (MSWD 1.0), which is distinctly
different from the mean age of the untreated group
when the 0.06 Ma internal error (based on analytical
precision only) of the latter is used for comparison.
A weighted mean 207 Pb/235 U age of 28.9 ± 1.5 Ma has
a very low MSWD of 0.20, indicating that the uncertainty on 207 Pb/ 204 Pb for the blank correction is
156
O. Bachmann et al. / Chemical Geology 236 (2007) 134–166
Fig. 9. Comparison of the zircon U–Pb data sets obtained in this study and by Schmitz and Bowring (2001). (a) and (b) show the ranked 206Pb/238U
ages of all main group zircons from both data sets, including their 1σ error, and (c) displays the corresponding age-probability curves. Whereas the
age distributions of both studies have similar upper limits, the data from the present study are skewed towards younger ages resulting in a slightly
younger weighted average. Open symbols in b) represent multi-grain analyses.
probably overestimated. Two additional CA-treated
crystals yielded consistent, but imprecise results due to
extremely small residual sample size. Fig. 10 shows
the concordia plot, with CA-treated and untreated
crystals represented by grey and white error ellipses,
respectively.
The analytical errors for most of the individual ages
measured at BGC are less precise on average than those
measured at ETH. This is partly due to the relatively
small grain size of the zircons available to BGC, and,
more importantly, to the fact, that the CA-pretreatment
applied to 5 out of 10 samples strongly attacked the
grains and left residues as small as 1.3 μg with significantly lower U and Pbrad concentration than the
untreated ones. Contrary to expectations, the PPD zircons
were extremely sensitive to the HF etching process. The
pretreated grains are similar to the ETH data set in that the
analyses have common lead contents in excess of the
analytical blank (a series of simultaneously processed
samples have yielded common lead contents close to
analytical blank levels). As no correlation between
sample size and common Pb content is observed, we
have chosen to correct common Pb using the composition
of the analytical background (see Table 7), the uncertainty of which encompasses the common Pb composition in feldspar reported by Riciputi et al. (1995;
206
Pb/ 204 Pb = 18.448, 207 Pb/ 204 Pb = 15.569, 208 Pb/
204
Pb = 37.665). Due to similar common Pb ratios in
FCT feldspar and our analytical blank, the reported ages
are relatively insensitive to variations in common Pb
content; using a combination of the feldspar values and
the analytical blank would bias the ages by an insignificant amount (+10–20 ka, see also below).
4.4. Discussion of the U–Pb data
4.4.1. Analytical and considerations and systematic
effects
Before the results can be interpreted in terms of
evolution of the Fish Canyon magmatic system, it is
important to establish whether the observed dispersion
of the 206 Pb/238U ages is analytically robust and thus
intrinsic to the analyzed zircon populations, or whether
it may relate to analytical artifacts and/or systematic
problems with the U–Pb dating method.
4.4.1.1. Influence of laboratory blank. The very small
amounts of radiogenic Pb present in many of the analyses
(20 samples have radiogenic Pb contents b10 pg) require
Fig. 10. U–Pb concordia representation for untreated and CA-treated
single zircons from the Pagosa Peak Dacite (PPDcc) analyzed at BGC.
O. Bachmann et al. / Chemical Geology 236 (2007) 134–166
tight control and quantitative treatment of analytical backgrounds; i.e., the amount of Pb in the blank, its isotopic
composition, and the attendant reproducibility, which vary
as a function of the amounts and isotopic composition of Pb
contributed from the various sources (loading blank, reagents, containers, exposure to the environment) are all
important. The multiple-source origin of our Pb blank led to
variably strong correlations between Pb isotopic ratios in
these blanks and weaker correlations between these ratios
and Pb content. We have included these correlations in the
data reduction and error calculation procedure detailed in
the Appendix. The required parameters have been obtained
by repeating blank experiments on both individual
analytical steps and full simulations of the zircon decomposition and U–Pb extraction path (a subset of these
parameters is shown in the caption of Table 6). As the
207
Pb/235U ages, which are very sensitive to analytical
blank, show good reproducibility, we consider the results to
be robust with respect to uncertainties arising from analytical contamination. This is also demonstrated by samples
such as GrdX2 34–37, which, although showing rather low
Pbrad/Pbcom (206Pb/204Pbb 100; Table 6), do not yield
aberrant U–Pb ages.
4.4.1.2. Isobaric interference and abundance sensitivity. Isobaric interferences in mass spectrometric
measurements are a further area of concern when
measuring radiogenic Pb in amounts as low as 0.9 pg
(e.g., GrdX2 37). In order to avoid contributions from
hydrocarbons, and thallium to the masses of interest,
Pb and UO2 were run at relatively high filament
temperatures of 1380–1450 °C and 1500–1550 °C,
respectively. During the Pb runs, mass 201 was
continuously monitored for BaPO2+ , which tends to
grow in that temperature range, in order to correct for
minor (typically insignificant) contributions to masses
204 and 205. Because barium is very efficiently
eliminated by the ion exchange procedure used for the
separation of Pb, such contributions are typically
negligible for zircon runs, but occasionally can cause
substantial shifts in Pb isotopic composition during
measurements of reagent blanks.
Because the single-grain samples have been spiked
with only ca. 10 pg 205Pb, their measured 205Pb/204Pb
ratios were ≤750. At an abundance sensitivity of b 2 ppm
(measured at mass 237), contributions from mass 205 to
mass 204 were b 0.15%. Such contributions have a negligible effect on the present data set.
4.4.1.3. Sample common lead correction. After corrections for procedural Pb blanks, most analyses still
contain a significant residual common lead component
157
that is related to the abundance of mineral and melt
inclusions in the analyzed zircons (Tables 6 and 7). To
correct for this contribution, we have chosen a mean
isotopic composition on the basis of four FCT feldspar
measurements reported by Riciputi et al. (1995;
206
Pb/204Pb=18.448, 207Pb/204Pb=15.569, 208Pb /204Pb=
37.665¸ see caption of Table 6). These values are
essentially identical to those of acid-leached FCT feldspars
measured by Schmitz and Bowring (2001) and suggest a
restricted range of common lead isotopic compositions for
the Fish Canyon magmatic suite.
In order to explore whether potential variations in
common lead composition might have resulted in an
artificial spread in 206 Pb/238 U ages, we consider samples FCTfv 12 and PPDcc 6, which have yielded the
highest (28.67 ± 0.13 Ma) and lowest ages (28.03±
0.09 Ma), respectively. For the 206 Pb/238 U ages of
these two samples to become identical to the weighted
mean age of 28.37 Ma derived from the total sample
population, the 206 Pb/ 204 Pb values for common lead
corrections on those two zircons would need to be
adjusted to 26.71 and 12.26, respectively. Both these
common Pb compositions are exotic in this geological
context. Similarly, in order to make zircons NCD 22,
GrnX 30 and GrdX1 33, which have young apparent
ages of 28.12 ± 0.08, 28.18 ± 0.09 and 28.16± 0.09 Ma,
respectively, match an age of 28.37 Ma, 206 Pb/204 Pb
values of 16.05, 17.40 and 17.42 would be required
for the respective sample common Pb components.
Such low values have neither been observed in the
central San Juan Mountain area (Lipman et al., 1978;
Riciputi et al., 1995) nor do they conform to the range
of measured procedural blanks (Tables 6 and 7).
4.4.1.4. Systematics of blank and sample common lead
correction. Our 206 Pb/238 U ages are relatively insensitive to uncertainties arising from fluctuations in the Pb
content of the analytical blanks, even if the ratio
of radiogenic to common Pb (and thus the measured
206
Pb/ 204 Pb ratio) is small. This can be verified by
examination of the following (simplified) equation for
radiogenic 206 Pbrad (=moles of radiogenic 206 Pb
generated in-situ in a sample)
206
Pbrad ¼ 206 Pbanalysis − 206 Pb blank −
2
6204
6
4 Pbanalysis −
ð
206 Pb
204
Pb
3
Pb blank 7
7;
5
206
206
ð Þ
204
Pb
Pb
blank
Þ
common
158
O. Bachmann et al. / Chemical Geology 236 (2007) 134–166
and its partial derivative of with respect to
(=moles of 206 Pb in the analytical blank)
206
Pbblank
206 Pb
204
A Pbrad
Pb
¼ 206 common −1;
206
Pb
A Pbblank
204
206
Pb blank
where the subscripts rad, analysis, blank and common
refer to radiogenic Pb, total Pb present in the analysis,
Pb in the analytical blank, and sample common Pb
(=inherent non-radiogenic Pb component of a sample),
respectively. The partial derivative tends to zero as the
two isotopic ratios (i.e., 206 Pb/ 204 Pb common and
206
Pb/ 204 Pbblank) converge. Therefore, only a fraction
of the (typically substantial) uncertainty associated with
206
Pbblank is actually propagated into the error of
206
Pbrad. If 206 Pb/ 204 Pb in an analysis is measured
with sufficient precision and 206 Pbblank does not
dominate the analysis, 206 Pbrad is thus determined to
considerably better precision than the 206 Pb content of
the sample, which suffers from the full uncertainty
associated with 206 Pbblank.
As 206 Pb/204Pb of the adopted sample common Pb
value lies in the middle of the range of measured
analytical blank isotopic compositions (see caption of
Table 6), our results are rather insensitive to variations
(and associated potential outliers) in blank lead
amounts. Whereas possible variations in sample common Pb isotopic composition and analytical uncertainties in sample/blank common Pb proportions could
influence our data at the 10 ka level, we conclude that
they cannot explain the total spread of the observed
206
Pb/238U ages.
We note that the four youngest samples (PPDcc 6,
NCD 22, GrnX 30, GrdX1 33) have relatively low
measured 206Pb/204Pb ratios (119–221; Table 6) as
compared to the average of main group zircons (255,
omitting GrdX2 37), resulting in a weak correlation
between 206 Pb/204Pb and age, which might point to an
analytical problem. However, there is also a weak
positive correlation between age and U concentration; in
such a case, 206Pb/204Pb and age are expected to become
correlated, as low-uranium samples have lower radiogenic 206Pb and consequently lower 206Pb/204Pb ratios.
Despite the below-average 206Pb/204 Pb ratios of these
youngest samples, their ages are shifted by less than
0.06 Ma if all common Pb is assumed to be analytical
contamination.
4.4.1.5. Initial radioactive disequilibrium. Zircon
crystallizing from magma preferentially incorporates
uranium relative to thorium, leading to initial radioactive 230 Th/238U disequilibrium in the zircon. Within a
few half-lives of the intermediate daughter nuclide
230
Th (T1/2 = 75.4 ka) 230 Th/ 238 U then approaches
secular equilibrium. The associated increment in 230 Th
is responsible for an equal deficit in 206Pb, the final
stable decay product of 238U. If 230 Th/238U in the melt at
the time of zircon crystallization was at secular equilibrium, the age offset caused by this effect can be approximated by Δ A g e ≈ (1 / λ 2 3 0 )( f − 1), where
f = (Th/U)zircon / (Th/U)melt, and 1/λ230 = 109 ka is the
mean life of 230Th (Barth et al., 1994). We have
approximated (Th/U)melt using the value of 2.2 measured
by Schmitz and Bowring (2001) on phenocryst-free pumice
shards from the FCT for correction of our 206Pb/238U ratios
and age data. This correction increments the 206Pb/238U
ages by 47 to 89 ka for (Th/U)zircon varying from 1.24
(PPD 6) to 0.38 (GrnX 27) for ETH analyzed zircons
and 65 to 80 ka (with Th/Uzircon from 0.87 to 0.52) for
those measured at BGC. Schmitz and Bowring (2001)
concluded that Th/U = 2.2 for FCT pumice most likely
represents evolved melt composition and thus should be
considered a minimum value for Th/U in the host
magma at the time(s) of zircon crystallization. This is
corroborated by higher Th/U measured in glasses
(2.34 – 2.71; Bachmann et al., 2005) and whole-rocks
(2.27–4.26, average: 3.03; Bachmann et al., 2002).
Adopting values of 3.0 or 4.0 for (Th/U)melt, the ages
would increase by trivial amounts (from + 0.005, or
+ 0.008, Ma for GrnX 27 to + 0.016, or + 0.028, Ma for
PPD 6, respectively) relative to corrections using a
value of 2.2. Thus, the uncertainties in our zircon ages
arising from incomplete knowledge of (Th/U)melt are
more than an order of magnitude smaller than the
maximum measured age differences.
4.4.2. Interpretation of zircon U–Pb results
The bulk of the main-group zircons analyzed from
the three main lithologies of the Fish Canyon magmatic
system, together with three out of four xenoliths collected from the FCT, yield ages that spread across a time
interval of 28.67 ± 0.13 Ma (FCTfv 12) or 28.61±
0.08 Ma (mean of CA treated PPDcc zircons) to 28.05 ±
0.09 Ma (PPDcc 6) (Figs. 8a–g, 9 and 10). In view of
the age scatter, the relatively limited number of U–Pb
analyses that we performed on six of the seven
lithologies that form the main-group data does not
permit us to establish significant differences among
these zircon populations; mean ages overlap within
errors and the populations are similar with respect to
their overlapping ranges of Th/U and (typically
moderate) uranium concentrations. An exception is
O. Bachmann et al. / Chemical Geology 236 (2007) 134–166
the youngest zircon analyzed in this study (PPDcc 6),
which has a higher Th/U ratio and a lower U concentration than the remaining main-group zircons.
Clearly distinct from the main group represented by
zircons with ages b 28.7 Ma are the three dates obtained
for a tonalitic xenolith from within the FCT, which give
a mean 206 Pb/238U age of 31.31 ± 0.05 (TonX 24–26,
Fig. 8j). This age makes the tonalitic xenolith a late
member of the 35–31 Ma magmatic cycle of intermediate composition lavas and breccias of the Conejos
Formation, which is mainly present around the perimeter of the San Juan volcanic field (Lipman et al., 1970).
One of the four zircons analyzed from a granitic xenolith
in the FCT (GrnX 27) shows an intermediate age of
30.29 ± 0.11 Ma, which is close to the eruption ages of
early silicic ash-flow sheets from the northern end of the
Central San Juan cluster (Cochetopa area; Lipman, personal communication). Zircon PPDcc 8, which hosted a
tiny bubble-core like structure in its center and was
suspected of potential inheritance, has yielded the only
analytically discordant result (Fig. 8h). A straight line
connecting the data point to the position of the mean age
of 28.369 Ma on concordia extrapolates to ca. 450 ±
120 Ma, most likely a minimum age for the inherited
component. Lanphere and Baadsgaard (2001) have documented, by conventional TIMS and ion probe methods,
substantial quantities of Precambrian zircon components
in their sample from the FCT.
Whereas there is clear separation between inherited
components with ages N 30 Ma and the b 28.7 Ma age
components, the substantial spread shown by the latter
poses a more complex problem. It raises the question of
whether the individual ages reflect a true (although
integrating) time scale for a magma chamber evolution
process lasting up to 0.6 Ma, or whether zircons crystallized at an early stage suffered variable degree of
secondary loss of radiogenic Pb at times pre-and/or
postdating the eruption events. We will examine this
question using the results obtained for the Pagosa Peak
Dacite (PPDcc; Tables 6 and 7; Figs. 8 and 10).
While the group of air-abraded samples and the
combined groups of untreated and CA-treated zircons show broadly overlapping age distributions, the
CA-treated zircons generally yield older U–Pb ages
(28.61 ± 0.08 Ma) than air-abraded (28.33 ± 0.15 Ma) or
bulk grains (28.46 ± 0.15 Ma; BGC data). Furthermore,
the CA-treated samples are clearly distinct from the
remainder by their low average uranium contents
(230 ppm) when compared to air-abraded zircons
(400 ppm) and bulk crystals (780 ppm). As in other
studies (Mundil et al., 2004; Mattinson, 2005), this
relationship suggests that younger ages are predomi-
159
nantly associated with U enriched domains, which
are more readily attacked and removed by the CA
procedure than low-U domains.
In this context, pre- and/or post-eruptive leakage of
some radiogenic Pb from narrow (μm or sub-μm thick),
U-enriched zircon domains could be invoked to explain the presence of younger age components in the
analyzed zircon populations, some of them possibly
even postdating eruption (if the true eruption age is
N 28.0 Ma). Pre-eruptive Pb loss by volume diffusion
(at temperatures N 700 °C; Bachmann et al., 2002) is
unlikely due to the fact that annealing is much faster
than accumulation of radiation damage at these temperatures (“critical amorphization temperatures” are as
low as ∼ 360–380 °C for U concentrations of 1000–
10000 ppm; Meldrum et al., 1998). At magmatic conditions, the diffusion parameters derived for nonmetamict zircon by Lee et al. (1997) and Cherniak and
Watson (2001) would preclude analytically resolvable
pre-eruptive Pb loss by volume diffusion in the given
amount of time (b ∼ 0.5 Ma), as those yield closure
temperatures in excess of 900 °C. However, in view of
petrological evidence for thermal oscillations in the
magma chamber prior to eruption, impure, U-enriched
zones in zircons may have undergone (multiple?) solidstate recrystallization events accompanied by Pb loss
prior to eruption (e.g., Schaltegger and Hoskin, 2003).
In contrast to crystallization/dissolution episodes affecting predominantly zircons exposed to interstitial
melt, such a solid-state reordering process would also
affect zircons enclosed in other minerals (which may be
a significant fraction of the total zircon population).
Pb loss occurring after eruption is also possible. The
glassy state of the sampled lithologies argues against
hydrothermal alteration of the rocks being responsible
for post-eruptive resetting of the U–Pb systems of the
FCT zircons, but Pb could have been lost by volume
diffusion from metamict domains. At low temperatures,
annealing rates are too low to fully compensate for
accumulation of radiation damage. Previous volume
diffusion experiments on zircon by ion implantation
carried out by Cherniak et al. (1991) have resulted in Pb
diffusivities orders of magnitude higher than those
obtained from the more recent determinations by Lee
et al. (1997) and Cherniak and Watson (2001). If the
high diffusivities observed in the ion implantation
experiments are taken as proxies for the diffusion
characteristics of natural metamict domains in zircon,
resetting of U–Pb ages by ∼ 0.5 Ma during a time
interval of 28 Ma at temperatures as low as ∼ 200 °C is
possible, given that the metamict zones would be
located at or close to the surface of the grains. However,
160
O. Bachmann et al. / Chemical Geology 236 (2007) 134–166
the fact that air abrasion was rather inefficient in
producing low U grains with older ages indicates that
the zones of U enrichment are not entirely confined to
the outermost shells of the crystals.
Partial resetting of the U–Pb system may have
occurred in FCT zircons, but is most likely not the only
mechanism responsible for the observed spread in U–Pb
ages. Assuming that batholithic magma bodies are built
incrementally (Wiebe and Collins, 1998; McNulty et al.,
2000; Mahan et al., 2003) over long time scales (e.g.,
Coulson et al., 2002; Sano et al., 2002; Coleman et al.,
2004), our preferred interpretation of the U–Pb results
presented in this study is episodic zircon crystallization
(perhaps accompanied by some minor recrystallization of
trace element-enriched zones) during an extended nearsolidus period with a duration of up to ∼0.3 Ma prior to
eruption of the FCT (allowing for an adjustment of the
Ar–Ar age to ∼28.3 Ma; see Section 4.4.4), with minor
overprint by post-eruptive open-system processes. This
protracted crystallization history at near-solidus conditions also agrees with: (1) petrological observations from
the Fish Canyon magma system and other large-scale
silicic magma bodies (Bachmann and Dungan, 2002;
Bachmann et al., 2002; Hildreth, 2004), (2) magma
residence times in excess of 200 ka estimated by U–Th–
Pb SIMS dating of zircons in silicic units smaller than
FCT (Reid et al., 1997; Brown and Fletcher, 1999;
Vazquez and Reid, 2002, 2004; Miller and Wooden, 2004;
Bacon and Lowenstern, 2005; Charlier et al., 2005;
Bachmann et al., in press), and (3) models of open-system
thermal evolution of large silicic magma bodies (e.g.,
Spera, 1980; Bachmann and Bergantz, 2003).
Such an interpretation requires that individual zircon
grains or growth zones originally crystallized at
different times (e.g., during cooling intervals following
pulses of heat input into the system) and/or in different
parts of the magma chamber (e.g., in its center or near
the roof) before being assembled shortly prior to
eruption. In this case, one would expect some chemical
variability between zircons. The Fish Canyon zircons
indeed show very complex textures and zoning patterns
(as most of the others mineral phases in the Fish Canyon
magma body do; Bachmann et al., 2002; Bachmann and
Dungan, 2002), in agreement with a prolonged
(∼ 0.3 Ma) crystallization history in a dynamic
environment undergoing thermal oscillations (see also
Miller and Wooden, 2004).
4.4.3. Comparison with the previous zircon study by
Schmitz and Bowring (2001)
Schmitz and Bowring (2001) performed 24 singlegrain and eight multi-grain measurements on FCT zircons
from sample FC-2 (distributed by the New Mexico
Geochronology Research Laboratory at New Mexico
Tech, also used as a 40Ar/39Ar standard). Twenty-three
single-grain and seven multi-grain determinations yielded
a disequilibrium-corrected precise weighted mean
206
Pb/238U age of 28.478 ± 0.024 Ma with MSWD =
0.97. The remaining two dates were clearly discordant and
thus indicative of inheritance. The age distribution
obtained is similar to that obtained in this study, wherein
a main group falls near a time interval centered around
28.3–28.6 Ma, and only few results manifest older
age components. In view of the pervasive presence of
inheritance in the results of Lanphere and Baadsgaard
(2001), these older components must either be heterogeneously distributed among the rocks analyzed, as the
samples originate from different locations, or they have
been very efficiently filtered by the crystal selection
procedures applied by Schmitz and Bowring (2001) and
in our study.
Whereas Schmitz and Bowring's (2001) data and
our study yield similar upper limits for the ranges of
disequilibrium-corrected 206 Pb/ 238 U ages (28.67±
0.13 Ma, zircon FCTfv 12, this study, and 28.62±
0.21 Ma, zircon z3 of Schmitz and Bowring (2001), the
youngest zircon age obtained in our study, 28.03 ±
0.09 Ma (PPDcc 6), is considerably younger than the
youngest age measured by Schmitz and Bowring (2001;
28.36 ± 0.12 Ma; z72b), disregarding z18, which has a
large error). The data patterns from the two studies are
compared in Fig. 9. The distinction is particularly
evident from the cumulative probability density plots
(Fig. 9c), wherein the distribution of our ages is skewed
towards younger dates, and by a shift in mean ages
(Fig. 9a and b). Considering only samples of FCT s.s.
(FCTfv and FCTar), the minimum age obtained by the
present study, 28.30 ± 0.11 Ma, is similar within error
limits to the minimum age determined by Schmitz and
Bowring (2001).
The narrow, Gaussian distribution displayed by their
206
Pb/238U dates, their MSWD value of 0.97 and the highprecision of their mean age led Schmitz and Bowring
(2001) to interpret their data in terms of a unique age for
the investigated zircon population showing no evidence
for magma residence time. Such an interpretation cannot
be applied to our data sets, which have MSWD values of
8.4 (ETH results) or 3.5 (BGC results for untreated
crystals) and thus clearly document age scatter. A review
of an earlier version of this paper proposed that the
analytical errors of our data may have been underestimated as compared those of Schmitz and Bowring
(2001), and it was recommended that we recalculate our
data by PbMacDat, a program applied by Schmitz and
O. Bachmann et al. / Chemical Geology 236 (2007) 134–166
Bowring (2001) to their data set (M. Schmitz, personal
communication 2005) rather than by our (previously
unpublished) data reduction procedure, which we now
show in the Appendix. The application of PbMacDat to
some of our data indeed resulted in errors of our
206
Pb/238U ages, which were approximately 3–4 times
larger than estimated by our program. An examination of
the code in PbMacDat prompted by this result has
revealed an error, which fully explains the discrepancy
(for details, see Appendix). If the data of Schmitz and
Bowring (2001) were indeed reduced with a similar
version of that program, their errors might have been
grossly overestimated. Consequently, by proper error
treatment, their data would become considerably more
precise than the published values. On the other hand,
correction for overestimation of analytical errors by a
factor of ∼3–4 would also raise MSWD by an order of
magnitude (since MSWD is approximately inversely
proportional to the square of the errors). MSWD ≈ 10 as
compared to the published value of 0.97 would then
clearly be at variance with the assignment of a unique age
to the analyzed zircon population and would be more
compatible with the results obtained in this study.
Independent of the above statistical arguments, the
narrowly focused age data set of Schmitz and Bowring
(2001), as compared to the more pronounced scatter in
our results, suggests differences in degree of age variations among the investigated lithologies and may also
reveal dependence of the age results on grain selection
and sample pretreatment (as demonstrated by the CA
technique).
4.4.4. Discrepancy between the 40Ar/39Ar and U–Pb
dating methods
The most commonly quoted 40 Ar/39 Ar age for the
Fish Canyon magmatic system is ∼ 28.0 Ma (Renne et
al., 1994; Renne and Min, 1998; Renne et al., 1998;
Villeneuve et al., 2000; Daze et al., 2003). If one accepts
a measured ∼ 28.0 Ma 40Ar/39 Ar date for the eruption
age of the Fish Canyon system and allows for an
increase by 1% related to potential bias in the decay
parameters of 40 K (Min et al., 2000; Renne et al., 2005),
then the resulting “corrected” eruption age of ca
28.3 Ma remains considerably younger than the oldest
zircon ages and the mean age for CA-treated crystals of
∼ 28.6 Ma found for the FCT. This strongly supports
our model for extended magma residence. Such a
conclusion seems unavoidable for a magma body as
large as the one that formed the Fish Canyon Tuff, in
light of the fact that several studies have now demonstrated that less voluminous magmatic systems had
zircon crystallization histories in excess of 200 ka. (e.g.,
161
Reid et al., 1997; Brown and Fletcher, 1999; Vazquez
and Reid, 2004; Bacon and Lowenstern, 2005, Charlier
et al., 2005; Bachmann et al., in press).
5. Conclusions
In contrast to the zircon results of Schmitz and
Bowring (2001), which have been interpreted as a
single age population with a 206 Pb/238 U age of 28.478 ±
0.024 Ma, our U–Pb dates on zircons of the Fish
Canyon magmatic system, collected from the most representative sample suite available, suggest an extended crystallization period spanning an age interval of at
least 0.3 Ma prior to eruption at ∼ 28.3 Ma (bias
corrected 40 Ar/ 39 Ar age on sanidine). This protracted
crystallization history is consistent with petrographic
and geochemical evidence for a long-lived magmatic
system that was maintained at temperatures above the
solidus (Bachmann and Dungan, 2002; Bachmann
et al., 2002). On the basis of the crystal-rich nature
of the magma at the time of eruption, complex zoning
patterns in feldspars and hornblendes (Bachmann and
Dungan, 2002), and thermal modeling (Bachmann and
Bergantz, 2003), we conclude that the magma was
stored as a crystal mush for most of its magmatic life,
undergoing local thermal oscillations (between ∼ 700
and ∼ 800 °C) that enabled zircons to grow periodically over an extended period of time. Although
incremental assembly of this huge volume of erupted,
homogeneous magma cannot be definitely proven, it
seems geologically highly unlikely that N5000 km 3 of
dacitic melt arrived suddenly at shallow level in the
crust and underwent one short cooling and crystallization cycle.
Even though mineral phases in the Fish Canyon
magma are conventionally inferred to be open to diffusive argon exchange until quenching during eruption,
our results (and those of Spell and McDougall, 2003),
manifest traces of inherited argon, which also can be
interpreted in terms of a protracted crystallization
history. Biotite, hornblende and plagioclase, which are
less readily reset than sanidine (McDougall and
Harrison, 1999), give mean ages that are somewhat
older than sanidine, and a 2-cm diameter feldspar
megacryst yielded a staircase Ar release spectrum
during an incremental-heating experiment (ages becoming older as the temperature is increased), a feature that
is observed when xenocrystic material is partially reset
after becoming engulfed in hot magma (Gillespie et al.,
1982).
The case for inherited argon and complex crystallization histories in large, silicic, crystal-rich, ash-flow
162
O. Bachmann et al. / Chemical Geology 236 (2007) 134–166
sheets is strengthened by both U–Th dating of allanite
(Vazquez and Reid, 2004) and 40 Ar/39 Ar results on
multiple mineral phases of the ∼ 74 ka Young Toba
Tuff. These results are very similar to those obtained
on the Fish Canyon system, as the 40 Ar/39 Ar sanidine
ages for the Young Toba Tuff are unequivocally
younger than those obtained on plagioclase, biotite,
and hornblende. This data set has been interpreted by
Gardner et al. (2002) and Thomas et al. (2003) as
recording xenocrystic contamination preceding the
eruption by only 10–100 years. This model has
serious limitations, the most crucial of these being the
problem of homogeneously distributing a large
quantity of xenocrystic material (most of the crystals
being euhedral at the time of eruption!) in 2800 km 3
of high-SiO2 rhyolite liquid in just a few years. We
favor the alternate hypothesis that plagioclase, biotite
and hornblende have the ability to preserve small
fractions of inherited argon in retentive sites (structural traps) over much longer time periods than
104 years.
Our interpretation that the Fish Canyon U–Pb and
40
Ar/ 39 Ar data record a protracted crystallization
history, which is in accord with textural features of the
Fish Canyon minerals, implies that the discrepancy
between the 40Ar/39Ar and U–Pb ages is not entirely
due to calibration errors in the 40Ar/39Ar method. Our
range of precise zircon U–Pb ages from ∼ 28.67 to
∼ 28.03 Ma with a large MSWD value (8.4) and the
mean age of 28.61 ± 0.08 Ma derived from CA-treated
zircons suggest that the difference between the older U–
Pb ages and the 40Ar/39Ar sanidine age is an expression
of extended magma assemblage, cooling, and rejuvenation and that a mean zircon age for the U–Pb data set
does not serve as a marker for intercalibration of these
methods.
Acknowledgments
This is part of a Ph.D. thesis by O.B. (Swiss FNRS
Grant # 20-49730.96 to M.D.). Many thanks to Brad
Singer, Yann Vincze and Thao Ton-That for advice
and help in the 40 Ar/39 Ar lab and to Peter Lipman for
his continuous support and invaluable guidance over
these last years. We are grateful to Irene IvanovBucher and Hannelore Derksen for carrying out part of
the mineral separation work required for this study.
Analytical work at the BGC is supported by the Ann
and Gordon Getty foundation. We thank Mary Reid,
Mark Schmitz and an anonymous reviewer for constructive reviews, which helped to improve the manuscript. An earlier version of the manuscript has also
benefited from a thorough and constructive review by
Samuel Bowring.
Appendix A. Estimation of errors of
and 207Pbrad/235 U
206
Pbrad/238U
We start from the following expressions for the
radiogenic 206Pb and 207Pb contents of a sample:
206r ¼ ½R65md ð1 þ FÞ−R65td 205t−206b−R64c ð1Þ
f½R46md R65md ð1−FÞ−R45t
205t−206bd R46bg
and
207r ¼ ½R76md R65md ð1 þ 2d FÞ−R75td 205t−206b
R76b−R74cd f½R46md R65md ð1−FÞ−R45t
205t−206bd R46bg
ð2Þ
where 206r, 207r = moles radiogenic 206 Pb, 207 Pb
in sample, R76m, R65m, R46m = measured 207 Pb/
206
Pb,206 Pb/205Pb, 204Pb/206Pb ratios, F = coefficient for
linear mass fractionation correction per amu, 205t = moles
205
Pb added with the tracer, R75t, R65t, R45t = 207
Pb/205Pb, 206Pb/205Pb, 204Pb/205Pb ratios of tracer,
206b = moles 206Pb in analytical blank, R76b, R46b = 207
Pb/206Pb, 204Pb/206Pb ratios in analytical blank, R74c,
R64c = 207Pb/204Pb, 206Pb/204Pb ratios of sample common Pb.
Using a Gaussian error propagation scheme, the analytical uncertainties of 206r (s206r) and 207r (s207r) are
estimated from
s2206r ¼
A206r
d sR65m
AR65m
þ
2 2
A206r
d sR46m
þ
AR46m
2 2
A206r
A206r
d sF þ
d s206b
AF
A206b
þ
A206r
d sR46b
AR46b
ð3Þ
2
covð206b; R46bÞ
þ2d
A206r A206r
d
A206b AR46b
O. Bachmann et al. / Chemical Geology 236 (2007) 134–166
and
s2207r
2 2
A207r
A207r
d sR76m þ
d sR65m
¼
ð4Þ
AR76m
AR65m
2 2
A207r
A207r
d sR46m þ
d sF
þ
AR46m
AF
2 2
A207r
A207r
d s206b þ
d sR76b
þ
A206b
AR76b
2
A207r
A207r A207r
d sR46m þ2d
d
þ
AR46b
A206b AR46b
ð
covð206b; R46bÞ þ
A207r A207r
d
A206b AR76b
covð206b; R76bÞ þ
A207r A207r
d
AR76b AR46b
Þ
covðR76b; R46bÞ ;
where the sx are the uncertainties of the respective variables
and cov(x,y) =sx·sy·ρ(x,y) denotes the covariance of pairs of
variables x and y with correlation coefficient ρ(x,y). The
partial derivatives of functions (1) and (2) evaluate to
A206r
¼ ½1 þ F−R64cd R46md ð1−FÞd 205t
AR65m
A206r
¼ −R64cd R65md ð1−FÞd 205t
AR46m
A206r
¼ ð1 þ R64cd R46mÞd R65md 205t
AF
A206r
¼ −1 þ R64cd R46b
A206b
A206r
¼ R64cd 206b
AR46b
and
A207r
¼ R65md ð1 þ 2d FÞd 205t
AR76m
A207r
¼ ½R76md ð1 þ 2d FÞ−R74cd R46md ð1−FÞd 205t
AR65m
A207r
¼ −R74cd R65md ð1−FÞd 205t
AR46m
A207r
¼ ð2d R76m þ R74cd R46mÞd R65md 205t
AF
A207r
¼ −R76b þ R74cd R46b
A206b
A207r
¼ −206b
AR76b
A207r
¼ R74cd 206b
AR46b
163
For the present context we assume that correlations
between the measured 207Pb/206Pb, 206 Pb/205Pb and
204
Pb/206Pb ratios can be neglected. Furthermore, the
isotopic ratios of sample common lead as well as the
tracer parameters are treated as constants rather than
random variables (see Section 4.4.2).
Due to the occurrence of common and correlated
variables in Eqs. (1) and (2), 206r and 207r are also
correlated. This is expressed by their covariance
covð206r; 207rÞ¼
A206r A207r 2
d
ds
ð5Þ
AR65m AR65m R65m
A206r A207r 2
þ
d
ds
AR46m AR46m R46m
A206r A207r 2
d
d sF
þ
AF
AF
A206r A207r 2
d
ds
þ
A206b A206b 206b
A206r A207r 2
þ
d
ds
AR46b AR46b R46b
A206r A207r A206r A207r
d
þ
d
þ
A206b AR46b AR46b A206b
A206r A207r
covð206b; R46bÞ þ
d
A206b AR76b
A206r A207r
d
covð206b; R76bÞ þ
AR46b AR76b
covðR46b; R76bÞ:
The uncertainties of 206 Pbrad/238U (=206r / 238U )
and 207Pbrad/235 U (=207r / 235U ) are then
qffiffiffiffiffiffiffiffiffiffiffiffiffiffiffiffiffiffiffiffiffiffiffiffiffiffiffiffiffiffiffiffiffiffiffiffiffiffiffiffiffiffiffiffiffiffiffiffiffiffiffiffiffiffiffiffi
1
s206r=238U ¼
s2206r þ ð206r=238U d s238U Þ2
238U
ð6Þ
and
qffiffiffiffiffiffiffiffiffiffiffiffiffiffiffiffiffiffiffiffiffiffiffiffiffiffiffiffiffiffiffiffiffiffiffiffiffiffiffiffiffiffiffiffiffiffiffiffiffiffiffiffiffiffiffiffiffiffiffiffi
1
s207r=235U ¼
s2207r þ d ð207r=235U d s235U Þ2 ;
235U
ð7Þ
where 238U, 235U, s238U and s235U refer to the molar
amounts and analytical uncertainties of 238U and 235U in
the sample as calculated from the double spike procedure,
with 235U = 238U / 137.88 and s235U = s238U / 137.88. The
covariance between 206 Pbrad/238U and 207Pbrad/235 U
evaluates to
1
covð206r=238U ; 207r=235U Þ ¼
ð238U Þ2
206r 207r 2
d
ds
137:88d covð206r; 207rÞ þ
238U 235U 238U
ð8Þ
164
O. Bachmann et al. / Chemical Geology 236 (2007) 134–166
The error propagation procedure adopted here is
similar to that of Ludwig (1980), apart from explicitly
including mass fractionation and omitting errors for
sample common lead composition. We also have slightly
varied input parameters for analytical blank (we prefer to
use 207Pb/206 Pb and 204Pb/206 Pb rather than 207 Pb/204Pb
and 206Pb/204 Pb in order to keep correlations between
these variables at a minimum) and have added correlations between amount of blank 206Pb and isotopic ratios.
The latter are frequently correlated due to varying blank
contributions originating from reservoirs characterized
by distinct isotopic composition.
The discrepancy between the analytical errors
calculated by program PbMacDat (downloaded May
2006 from the http://www.earth-time.org website),
which implements the error estimation procedure of
Ludwig (1980), and by the procedure described here
(see discussion in Section 4.4.3), has been traced to an
error
term
P2 in PbMacDat, where the original (correct) P
2
in
Ludwig's
(1980)
Eq.
(19)
is
replaced
by
CT P
CS ,
P2
2
with CT and CS denoting the fractional variances of
total 206Pb in the analysis (CT) and of sample 206 Pb
(CS), respectively. Because the latter expression carries
the full weight of the (typically substantial) uncertainty
in the amount of blank to be subtracted from an analysis
and lacks the additional control by 204Pb (see Section
4.4.2), the analytical errors become considerably
inflated.
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