Chemical Geology 257 (2008) 114–128
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Chemical Geology
j o u r n a l h o m e p a g e : w w w. e l s ev i e r. c o m / l o c a t e / c h e m g e o
Determining high precision, in situ, oxygen isotope ratios with a SHRIMP II: Analyses
of MPI-DING silicate-glass reference materials and zircon from contrasting granites
R.B. Ickert ⁎, J. Hiess, I.S. Williams, P. Holden, T.R. Ireland, P. Lanc, N. Schram, J.J. Foster, S.W. Clement
Research School of Earth Sciences, The Australian National University, Canberra, ACT, 0200, Australia
a r t i c l e
i n f o
Article history:
Received 26 June 2008
Received in revised form 21 August 2008
Accepted 24 August 2008
Editor: R.L. Rudnick
Keywords:
SIMS
SHRIMP
Zircon
Oxygen isotopes
Granite
a b s t r a c t
The development of new techniques and instrumentation on the ANU SHRIMP II ion microprobe has made it
possible to measure the oxygen isotope ratios of insulating and conducting phases (e.g. silicates, carbonates,
phosphates and oxides) on a 25 µm scale with better than 0.4‰ precision and accuracy at 95% confidence.
Instrumentation changes include the installation of a multiple collector, charge neutralization using an
oblique-incidence low-energy electron gun, and the addition of Helmholtz coils to counter mass dispersion
by the Earth's magnetic field. A redesign of sample mounts and mount holders has effectively eliminated
differences in variable isotope fractionation across the mount surface during analysis. Techniques have been
developed to minimize the effect of electron-induced secondary ionization of oxygen. During a 6-minute
analysis involving 100–140 s of data collection, δ18O values can be measured on one 25 µm spot with an
internal precision of better than 0.2‰ (2 standard errors). Analyses of MPI-DING silicate-glass reference
material demonstrate that the external reproducibility of single spots can be better than 0.4‰ at 95%
confidence, and that for matrix-matched samples and reference material, accuracy is commensurate with
precision. MPI-DING glasses are acceptable ion microprobe reference materials for oxygen isotope
measurements of glasses, although KL2-G is possibly heterogeneous. Zircon reference materials TEMORA 2
and FC1 appear to be acceptable as preliminary oxygen isotope reference materials. SHRIMP II analyses of
FC1 indicate that it has a δ18O value of 5.4‰ (VSMOW). Analyses of zircon oxygen isotopic compositions
from a gabbro, a tonalite and a granodiorite from southeastern Australia are presented. Zircon from the
gabbro has a δ18O value of 5.6‰, the tonalite has an I-type affinity and slightly heterogeneous δ18O values
around 6.6‰, and the granodiorite has an S-type affinity and a range of igneous, melt precipitated zircon
δ18O values between 8.2 and 10.2‰. These results suggest that the gabbro is mantle-derived and slightly
contaminated with crustal material, and that the I-type granodiorite has evolved in a similar manner from
a mantle-derived source. The δ18O values of the zircon from the S-type granodiorite are not only higher
than from the I-type, but also more heterogeneous, consistent with partial melting of a poorly-mixed,
metasedimentary source.
© 2008 Elsevier B.V. All rights reserved.
1. Introduction
SHRIMP (Sensitive High Resolution Ion MicroProbe) instruments
have been used in a wide variety of applications in the Earth sciences.
Principal among these is U–Pb geochronology because of the presence
of coherent age domains in the mineral zircon. SHRIMP has also been
effective in isotope cosmochemistry and trace element abundance
measurements. One of the first types of measurement carried out on
SHRIMP was the analysis of stable S isotopes. Initially this was
attempted with an Ar+ primary beam and negative secondary ions (S−),
⁎ Corresponding author. RSES, Building 61, Mills Road, Acton, ACT, 0200, Australia.
Tel.: +61 2 61250316; fax: +61 2 61250738.
E-mail address: [email protected] (R.B. Ickert).
0009-2541/$ – see front matter © 2008 Elsevier B.V. All rights reserved.
doi:10.1016/j.chemgeo.2008.08.024
although it was noted by Coles et al. (1981) that Cs+ primary ions would
produce much higher secondary ion yields. The success of the
instrument for the analysis of positive ions, however, and its ability
to make per mill measurements of S isotopes in that mode as well,
stalled further development of negative secondary ion isotope ratio
measurements. We have now revived negative ion work because of the
clear benefits of making multiple types of analyses, including oxygen
isotope ratio measurements, on single zircon domains (e.g., Mojzsis
et al., 2001; Peck et al., 2001).
A wide range of geological processes can modify the relative
abundances of oxygen isotopes, making oxygen isotope ratios a valuable
tracer in the geological and environmental sciences (Hoefs, 2004). Such
processes include equilibrium isotopic exchange, and incomplete or
unidirectional effects such as diffusion, evaporation/condensation, and
metabolism. Geological applications of oxygen isotope analysis include studies of magma genesis, hydrothermal systems, ore genesis and
R.B. Ickert et al. / Chemical Geology 257 (2008) 114–128
paleoclimatology (Sharp, 2007). Oxygen in extraterrestrial materials has
provided key information about the origin of the solar system and
nucleosynthesis (MacPherson et al., 2008). Oxygen isotope ratios of
zircon have proved particularly valuable in studies of magma genesis
because zircon can preserve the original oxygen isotopic composition of
the melt from which it precipitated, even if the subsequent igneous rock
has been heavily altered (Valley et al., 1994; Valley, 2003).
Secondary Ion Mass Spectrometry (SIMS) is a technique that can be
used to measure oxygen isotope ratios of solid samples in situ on the
micrometre scale (Ireland, 2004 and references therein). SIMS works by
focusing a beam of high-energy ions onto the solid under investigation,
resulting in the ejection or “sputtering” of material of interest. The
sputtered products include neutral atoms, molecular fragments, and a
small fraction of ionized particles that can be introduced into a mass
spectrometer, selected for energy and mass and measured by an ion
counter or Faraday cup. Oxygen is strongly electronegative and therefore
best measured as a negative ion, with strongly electropositive Cs+ as the
most common choice for primary ion species. When measuring
insulating samples, the configuration of positive primary ions and
negative secondary ions causes a build-up of a positive charge on the
sample surface, despite a conductive coat applied to that surface. For in
situ work, the charging problem is generally avoided by delivering
electrons to the sample surface (Migeon et al., 1990; Hervig et al., 1992),
although in special cases, small samples can be pressed into soft metals
such as In or Au (McKeegan, 1987).
Early oxygen isotope ratio measurement by SIMS instruments were
limited by count times on the order of 30–45 min, and a precision of
about 1‰ (1σ; McKeegan, 1987; Valley and Graham, 1991; Hervig et al.,
1992; Graham et al.,1996). These limitations were primarily a function of
the inherently low transmission of the small turning radius SIMS (such
as the Cameca IMS f series), the need for extreme energy filtering, and
the measurement of oxygen isotope ratios dynamically on a single
electron multiplier. The development of the large turning radius Cameca
IMS1270 for oxygen isotope analysis (de Chambost et al., 1994;
McKeegan et al., 1998) overcame these obstacles, providing both higher
transmission and dispersion while generating stronger secondary beam
currents (higher count rates) and permitting the use of static multiple
collection (de Chambost et al., 1998). These instruments have achieved
improved precision and reproducibility for substantially faster analyses.
Many of the developments necessary for these measurements were
specific to the Cameca ion optical configuration, for example, use of an
electron gun with normal incidence in order to overcome the high field
gradient in the secondary ion extraction region.
115
Here we report major improvements to both the analytical
techniques and the instrumentation on the ANU SHRIMP II that
permit the high precision measurement of δ18O. Oxygen isotope ratio
measurements of MPI-DING silicate-glass reference material and
zircons used as U–Pb reference material (TEMORA 2 and FC1) demonstrate that better than 0.4‰ (at 95% confidence) external reproducibility can be attained. As an example, we have applied these
techniques to measure the δ18O values in zircon from two granites and
a gabbro from the Lachlan Fold Belt, southeastern Australia.
2. Configuration of the ANU SHRIMP II for oxygen isotopic analysis
The ANU SHRIMP II is a high mass resolution, double-focusing mass
spectrometer which is based on an ion optic design by Matsuda
(1974). The turning radii of the electrostatic and magnetic sectors are
1272 and 1000 mm respectively (Fig. 1). Although a secondary ion
energy window is available on SHRIMP II, it is generally not used (i.e. is
wide open) in isotope measurements because of the high quality of
the chromatic refocus. SHRIMP II incorporates an astigmatic secondary
ion beam-matching system to improve transmission through the
source slit and mass analyzer. The geometry and ion extraction system
are optimized for both high mass resolution and sensitivity. The
SHRIMP II was commissioned in 1992 as a single collector instrument
with a duoplasmatron primary ion source (Clement and Compston,
1990; Ireland et al., 2008). Both a Cs+ source and multiple collection
were envisaged for this instrument, with construction of a dual ion
source and multiple collector soon thereafter.
In a similar fashion to the development of a stable isotope analysis
capability on the Cameca instruments, it has been necessary to install
additional components on SHRIMP II in order to measure oxygen
isotope ratios with high precision. These include a demountable Cs+
primary ion source, electron gun, multiple collector, and Helmholtz
coils around the source chamber. The configuration of this hardware is
specific to the SHRIMP II geometry and is discussed below.
2.1. Cs+ primary ion source
The ANU SHRIMP II is equipped with a Kimball Physics IGS-4 alkali
metal ion gun fitted with a Cs+ firing unit. The gun can be removed and
stored in air when not in use. Ions are produced by thermal ionization
by heating a zeolite solid. The primary beam current is adjusted by
changing the temperature of the ion source. The gun is mounted on
an adaptor flange that makes it directly interchangeable with the
Fig. 1. Schematic diagram of the main elements of the ANU SHRIMP II. Key components discussed in the text are labelled.
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R.B. Ickert et al. / Chemical Geology 257 (2008) 114–128
Fig. 2. Schematic diagram of the main elements of the secondary ion extraction system,
electron gun, and Cs+ source. The electron beam is deflected by the difference in
potential between the sample and the intermediate extraction electrode, and is steered
to be coincident with the impact point of the Cs+ primary beam. Secondary ions are
initially accelerated to 0.7 kV by the intermediate extraction lens held 2 mm from the
sample surface. Full acceleration to 10 kV occurs after the intermediate extraction lens.
After Ireland (2004).
duoplasmatron ion source used for generating the oxygen (O−, O−2)
primary ions used in geochronology and trace element analysis. The
adaptor places the point of Cs+ emission in the same effective ion optic
position as the point of O ion emission from the duoplasmatron. The
source assembly is operated at close to ground potential, 10 kV above
sample potential. Caesium ions are extracted from the gun at an
additional 5 kV, giving the Cs+ ions a total energy of 15 keV at the
target (Fig. 2).
Lenses and apertures in the primary column demagnify and
collimate the Cs+ beam before projecting a sharp-edged spot of
uniform density onto the sample surface at an angle of 45°. First, a pair
of einzel lenses (one inside the Cs+ source assembly and the other part
of the “normal” primary column) works as a zoom-pair and project a
magnified image of the object formed by the Cs+ source onto the
differential pumping aperture at the entrance to the source chamber.
An ion image of that aperture is then projected forward by two einzel
lenses that have an aperture (the Köhler aperture) inserted between
them at the focal point of the second lens. In this form of Köhler
illumination, the primary ion beam is focused to an evenly illuminated
spot that is a demagnified image of the Köhler aperture (Ireland, 1995;
Williams, 1998).
During a typical analysis, a 3.5 nA beam of Cs+ is focused into a spot
∼25 µm diameter on the target surface, generating approximately
250 pA of O− secondary ions from most silicate, phosphate and
carbonate targets. The secondary ions are initially accelerated away by
a low extraction voltage (−700 V). The full acceleration to true ground
potential (i.e. +9.3 kV relative to the sample) is produced by an
extraction cone behind the extraction plate (Fig. 2).
2.2. Electron gun
The ANU SHRIMP II is equipped with a Kimball Physics ELG-5
electron gun. It is mounted on the side of the secondary extraction
lens assembly in the source chamber, with the nose of the gun 20 mm
Fig. 3. Schematic diagram showing the main components of the ANU SHRIMP II multiple collector.
R.B. Ickert et al. / Chemical Geology 257 (2008) 114–128
117
from the sample surface at an incidence angle of 45° (i.e., 90° from the
primary beam trajectory). Both the gun and its power supply are
floated at column ground, nominally −10 kV relative to true ground
(Fig. 2).
The electron gun uses a refractory metal cathode to generate an
electron beam with low-energy spread. Adjusting the cathode
temperature controls the beam intensity. The electron emission
energy is adjustable up to 3 keV. Under typical analytical conditions,
the gun provides ∼1 µA of electrons to the target surface. The gun
assembly includes two focusing elements that together operate as a
zoom lens, plus two pairs of deflection plates to control the beam
trajectory. The latter are particularly important as the electron beam is
strongly deflected as it enters the secondary ion extraction field
(∼ 350 V/mm). When focused, the electron beam, as imaged on a
phosphorescent plate, is concentrated in an ellipse approximately
100 × 200 µm. In practice, the electron energy is normally set in the
range 1.1–1.3 keV which, because of the ∼700 V secondary ion
extraction potential, illuminates the analyzed area with electron
energies in the range 400–600 eV. At such energies the electrons have
no observable effect on the conducting coating on the sample, or on
the epoxy mounting medium (cf. Hervig et al., 1992). Electron energy
is kept constant during an analytical session.
The electron beam is adjusted to provide the highest intensity
secondary ion current. Focusing and steering are very stable and the
gun can be left on while moving the sample. Due to the stability of the
charge neutralization, it is often the case that the electron gun does
not need to be adjusted during an analytical session, and after a
sample change only a small change in steering is required to reoptimize charge neutralization.
2.3. Multiple collector
The SHRIMP II multiple collector (Ireland et al., 2008) has five
collectors, three in a central array adjustable for single-mass-unit
spacing over the mass range 176–208, and two moving heads on either
side of the central array, with a maximum mass separation of 1 in 8
(Fig. 3). The two moving heads and the axial array are each equipped
with a single Faraday cup, linked by a low-capacitance high-vacuum
feed-through to an external Keithley 642 low-noise electrometer. The
moving heads are used for 18O/16O analysis. The electrometers
measuring 16O− and 18O− have 1010 Ω and 1011 Ω input resistors,
respectively. When necessary, the Faraday cups can be interchanged
with Sjuts sintered ceramic channel electron multipliers without
breaking vacuum.
The mass resolution required for measuring 18O/16O is relatively
low. Potential isobaric interferences on 18O from 17OH and 16OD can be
resolved with resolutions of 2300 and 1830 respectively. This is
achieved with 300 µm collector slits and a 150 µm source slit, which
truncates the secondary ion beam by less than 5%. Representative
peak shapes are shown in Fig. 4. An array of four beam-defining slits of
different widths is provided for each collector (100, 200, 300, 400 µm
on the moveable heads, and 50, 100, 200, and 300 µm on the central
array), allowing mass resolution to be adjusted in steps from ∼1500 to
N7000.
2.4. Helmholtz coils
The ambient magnetic field of the Earth is strong enough to
generate significant mass dispersion of oxygen isotopes between the
sample and the source slit, and can induce a component of
instrumental mass fractionation (IMF) that is sensitive to secondary
ion steering. After extraction and acceleration from the sample
surface, the secondary ions are focused by a triplet of quadrupole
lenses before passing through the source slit and entering the
secondary mass analyzer. Over the 350 mm from the sample surface
to the source slit, the ions are subject to the ambient magnetic field of
Fig. 4. Simultaneous mass scans of 16O and 18O showing flat-topped peaks and lack of
interferences. The vertical axes are linear in (a) and (b), and logarithmic in (c).
the Earth, the vertical component of which is ∼0.53 G in Canberra.
This field generates mass dispersion, with 16O− being deflected about
3.5 µm more than 18O−. As the edges of the secondary ion beam are
truncated by the source slit, this introduces a component of IMF that is
very sensitive to the horizontal steering of the beam. A similar effect
has been reported on other large radius ion microprobes (e.g.,
Schuhmacher et al., 2004).
The installation of coupled horizontal Helmholtz coils above and
below the source chamber has enabled the cancellation of the
terrestrial magnetic field and effected a substantial reduction of IMF
at the source slit of SHRIMP II. The IMF is relatively insensitive to the
vertical steering; no compensation is required in that direction. The
effectiveness of the coils in reducing steering-dependent isotope
fractionation is illustrated in Fig. 5. Without coils, the fractionation of
measured 18O/16O caused by changing the secondary beam horizontal
steering enough to truncate the beam by about 10% is 58‰ for a 50 µm
source slit, falling to 32‰ for a 150 µm slit. At a coil current of ∼ 0.38 A
(corresponding to a magnetic field strength of ∼ 0.5 G), the fractionation for the same deflection for all slit widths falls below detection
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the very centre of the mount. Only within 4 mm of the mount edge,
where the mount did not fully cover the 8 mm diameter extraction
aperture, did significant fractionation become apparent.
2.6. Electron-induced secondary ion emission (EISIE)
Fig. 5. Drift of secondary ions from the sample to the source slit through the Earth's
magnetic field results in mass dispersion, which is manifested by a change in the
measured isotope ratio as a function of ion beam steering against the source slit. The
magnitude of the difference in measured isotope ratio between a secondary beam that
is directed at the centre of the source slit and a secondary beam that is slightly offset is a
function of the magnitude of the mass dispersion. This difference (for three different
sized slits) is plotted vs. the current in the Helmholtz coils, which produce a different
strength magnetic field according to coil current. The difference in measured isotope
ratio (here depicted as δ18O values) between a secondary beam going directly through
the slit, and one that is slightly offset, is zero at a coil current of 0.4 A. This is the current
at which the Helmholtz coils produce a magnetic field that completely compensates for
the Earth's magnetic field.
When using the SHRIMP II electron gun during the measurement of
insulating materials, there is often a small but measurable O− signal
that can be detected at the multicollector when the Cs+ beam is
deflected away from the target, the Cs+ source power supply is off, or
the Cs+ source is physically isolated (by a gate valve) from the target.
This signal is not present when measuring oxygen isotope ratios in
conducting materials and the electron gun is not used, nor is it present
when the electron gun is running on a completely fresh, insulating
target before the Cs+ source has been turned on. The intensity of the O−
signal is positively correlated with the proximity of the electron
(b1‰). As the coil current is increased further, the fractionation
increases, but in the opposite sense. By combining the use of
Helmholtz coils with a 150 µm source slit, variable IMF as a function
of horizontal beam steering has been effectively eliminated. This has
desensitized the IMF to small adjustments in secondary beam steering
that are still required to compensate for microtopography on the
target surface and small differences in positioning the target plane.
2.5. Redesign of sample mounts
The standard SHRIMP II sample mount consists of a 25 mm
diameter epoxy disc supported by a stainless steel mount holder. The
face of the disc sits flush behind a thin (250 µm) flange in the mount
holder, so that the surface presented to the face of the extraction lens
is a 22 mm diameter Au-coated polished epoxy surface, surrounded by
a 6 mm wide stainless steel annulus, with a 250 µm step in between.
The IMF of oxygen secondary ions across even the central 10 mm
square of this sample mount was found to be large, with fractionations
in 18O/16O relative to the mount centre ranging from −4.5‰ at the NE
corner to +2.9‰ at the NW corner (Fig. 6). Within 2 mm of the inner
edge of the stainless steel annulus the fractionations were extreme,
−48‰ in the E and +22‰ in the W (not shown in Fig. 6).
To test whether the fractionations were related to the 250 µm step
near the mount edge, the presence of the stainless steel, or both, a
custom mount was prepared with a recessed edge such that the epoxy
and stainless steel flange defined a flat plane, eliminating the 250 µm
step. The result was to remove the extreme edge effects close to the
flange and the E–W gradient in relative mass fractionation, but the
fractionation gradient from centre to edge remained (about +4‰,
Fig. 7). This showed that the step in the mount surface produced the
pronounced mount edge effect, but that the contrast in material
between Au-coated epoxy and stainless steel also was a major
contributing factor and would still be a problem for SHRIMP II analysis.
To eliminate both the step and the material contrast between the
sample mount and its holder, the epoxy mount was redesigned to be
larger (35 mm diameter) and to attach to the face of the mount holder.
This “megamount” presented a uniform surface to the extraction lens
with no internal boundaries. The result was to effectively eliminate
the isotopic gradient over the central 19 mm (Fig. 8; Table 1). Isotopic
variation at any point in this area is the same as what is expected at
Fig. 6. Conventional mount holders for epoxy grain mounts present a material and
topographic contrast to the extraction field. This introduces a gradient in instrumental
mass fractionation across the mount and extreme, asymmetric edge effects near the
steel. Black diamonds are zircon grains with a size exaggerated for clarity. In the cross
section, the distance by which the steel sits proud of the mount face is exaggerated for
clarity.
R.B. Ickert et al. / Chemical Geology 257 (2008) 114–128
119
Electron-induced secondary ion emission of oxygen is observed on
SHRIMP II for a wide range of insulating target materials and operating
conditions. It effectively results in the simultaneous measurement of
two sources of oxygen: one directly sputtered from the spot by
primary Cs+, and a second produced by desorption over a wider area.
The EISIE is temporally variable, particularly during the sputtering
Fig. 7. An experimental mount holder was designed to present a flat face to the
extraction field where the epoxy is flushed with the steel. This design eliminated the
extreme edge effects of the conventional mount holder, but the isotopic gradient across
the face of the mount remained.
bombardment area to any sputtered crater (Fig. 9), and with electron
energy. There is no sample damage (e.g., removal of conductive coat)
associated with this effect.
We suspect that this anomalous secondary signal is the result of
electron-stimulated desorption—ESD (Madey, 1986). We refer to the
specific effect seen on SHRIMP II as electron-induced secondary ion
emission (EISIE). ESD is the production of low-energy ions and
neutrals through the electronic excitation of terminal bulk atoms or
adsorbed monolayers by electron bombardment (Madey, 1986). It is
commonly observed when an electron gun is used to compensate for
charging during SIMS analysis of electrical insulators (Williams, 1981;
Williams et al., 1983; Williams and Gillen, 1987). ESD can be a serious
problem, for example inhibiting the analysis of fluorine (e.g. McPhail
et al., 1986) and carbon (e.g. de la Mata and Dowsett, 2007), or a useful
tool, for example providing a means of visualizing the electron beam
when adjusting an electron gun (e.g. Reger et al., 1997).
Fig. 8. The newly designed “megamount” attached to the front of a modified sample
holder eliminates topographic and material contrast across the analytical surface. Mass
fractionation associated with position on the mount is only significant near the edge,
where the extraction aperture overlaps with the edge of the mount.
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Table 1
Analyses of zircon in transects across the face of a megamount
Spota
δ18Ob
X1
X2
X3
X4
X5
X6
X7
X8
X9
X10
X11
X12
X13
X14
X15
Y1
Y2
Y3
Y4
Y5
Y6
Y7
Y8
Y9
Y10
Y11
Y12
7.9
8.0
7.7
7.6
7.8
7.8
8.2
7.7
8.2
8.3
8.3
7.8
7.7
8.0
34.5
8.3
9.2
8.9
8.6
8.7
8.4
8.2
8.6
8.7
8.9
7.6
7.9
a
Numbers represent horizontal (X) positions from left to right on the mount, and
vertical (Y) positions from top to bottom.
b
Data are normalized to a mean δ18O value of 8.2‰. δ18O (‰) = ((18O/16OSample) / (18O/
16
OVSMOW )− 1) ⁎ 1000 18O/16OVSMOW = 0.0020052 (Baertschi, 1976).
process when it increases approximately and exponentially with time
(Fig. 10). The desorbed oxygen is isotopically fractionated by
approximately −200‰ relative to the Cs+ sputtered oxygen and has
a much lower mean energy and energy range.
Despite the low intensity of the EISIE—usually it contributes about
0.1% of the total measured 16O ions—its fractionated isotopic
composition and temporal and spatial variability mean that it has
the potential to affect analyses of target oxygen isotope ratios. We
have found that this effect is substantially reduced by using an Al
Fig. 9. Electron-induced secondary ion emission (EISIE) occurs near a sputtered region,
but dies away to a “background” EISIE within a few hundred micrometres. The data for
this figure was collected by sputtering the zircon grain for several minutes, then turning
the Cs+ beam off, moving the stage to different positions around the sputtered area and
measuring the count rate of 16O− due only to electron-induced ionization. The contours
were drawn by hand around the original spot analyses, and are labelled in per cent
relative to the maximum EISIE count rate determined near the analytical spot.
Fig. 10. Electron-induced secondary ion emission (EISIE) occurs when the electron beam
ionizes adsorbed oxygen on the sample surface. The measured abundance of these ions
increases approximately and exponentially with time, although it is low during the
length of a typical analysis. The data here were collected by sputtering a zircon under
normal analytical conditions, and periodically turning off the Cs+ beam and measuring
only the 16O ions desorbed by the electron beam. Time zero is when sputtering began.
conductive coat, rather than the Au coat that is used for most SHRIMP
II work, and by keeping the electron energy as low as possible while
still maintaining charge neutralization. In addition, analytical strategies have been adopted such that spots in both unknowns and
standards, where possible, are positioned so that the relative spacing
is similar or identical so that there is no bias between sets of analyses.
Finally, the EISIE intensity is usually measured 3 times during an
analysis, (by turning off the Cs+ and measuring the remaining 16O−
intensity) providing a check that the effect has been properly
minimized and data from which peak-stripping can be carried out,
although this is commonly not required (see Section 2.9 and 3.2 for
examples).
2.7. Analyses
Under the analytical conditions described above, the Cs+ beam
removes the ∼12 nm Al coating at the analytical site within a few
seconds and the secondary ion current rises steadily for about 3 min,
after which the rise slows or the beam steadies and then slowly declines,
depending on the matrix. During the rapid rise phase, the 18O−/16O− also
rises as steady state conditions between the isotopic composition of the
analyzed area, secondary ion beam, and implanted Cs become
established (Fig. 11). Before the secondary beam stabilizes, the primary
beam is rastered across an area slightly larger than the analytical pit for
30 s in order to remove any surface contamination. The spot is then
sputtered for up to an additional 90 s without rastering. After the
secondary beam stabilizes, the secondary ion steering is optimized,
the secondary ion beams are centred in the collector slits and the
data collection commences. Isotope ratios are normally measured for
100–140 s as one set of ten, or two sets of five to seven, 10-second
measurements, with re-optimization of the beam centring on the source
and collector slits between the two sets. The EISIE is monitored
throughout, with measurements taking place: 1) a few seconds after
the primary beam is turned on, 2) before data collection starts, and
3) after the isotope ratio data have been collected.
Typical count rates on 18O− are about 4–6·106 cps, and on 16O−
about 2–3·109 cps. Background count rates for 18O− and 16O− are
about 3.5·103 cps (1011 Ω resistor) and 11.5·103 cps (1010 Ω resistor)
respectively, and are measured at the start and end of each analytical
session.
2.8. IMF and gain correction
Due to the large IMF in SIMS, accurate isotope ratio measurement
requires standardization to a matrix-matched reference material of
R.B. Ickert et al. / Chemical Geology 257 (2008) 114–128
Fig. 11. Plot of the time-dependent rise in 18O/16O following first-impact of a ∼3 nA Cs+
ion beam on an Al-coated grain of TEMORA-2 zircon. Data collection commences once
the signal has stabilized, normally after about 3 min. Each data point represents a count
time of 10 s, and error bars are at 1SD and are based on counting statistics. More scans
are shown here than taken for a normal analysis, to demonstrate the negligible change
in the ratio after about 200 s.
121
Fourteen spots were analyzed on each of the six glasses, for a total
of 84 individual oxygen isotope analyses (Table 1). Oxygen isotope
ratios are reported in terms of the IMF, as the per mill deviation of the
gain and background corrected 18O/16O from the true 18O/16O (e.g.,
Eiler et al., 1997). Because the difference between the true isotopic
compositions of the materials and VSMOW is small (a maximum of
9.4‰), the size of a 1‰ variation in IMF is nearly identical to a 1‰
variation in a δ18O value relative to VSMOW. Each analysis took less
than 6 min, and the results are illustrated in Fig. 12. Ten scans, of 10 s
each, were measured for each analysis, and count rates on 16O ranged
from 2.5 to 2.9·109 cps. Electron-induced secondary ions accounted for
approximately 0.03% of the total secondary beam, and induced very
small changes in the measured isotopic ratios, on the order of 0.1‰.
The change is systematic for all spots. For each spot, the additional
uncertainty due to fluctuations in the EISIE is very small and on the
order of 0.01‰; as such, no correction was applied.
The precision of an individual analysis (i.e. an individual spot) can
be described by three different values in an hierarchical manner (e.g.,
Stern and Amelin, 2003); in decreasing order of precision they are,
“internal precision”, “spot-to-spot precision”, and “external precision.” For the following discussion, we define the standard deviation,
or dispersion of the data, as SD, and the standard error (or standard
deviation of the mean, SD/√N, where N is the number of samples), or
known and nominally homogeneous isotopic composition. During an
analytical session, analyses of the reference material are interspersed
with those of unknown samples. The analyses of reference material
serve to determine the magnitude of the IMF, which can change on a
day-to-day basis, and to monitor any changes during an analytical
session. Normal practice is to measure the reference material
(standard) several times at the beginning of an analytical session,
and thereafter following every three to five sample analyses.
Both IMF and relative gain between electrometers are corrected by
normalizing all background corrected data to the measured isotopic
composition of the reference material. For single collector data,
following Eiler et al. (1997), the IMF can be depicted as an α ratio
Measured
α u RTrue
Std =RStd
where R = 18O/16O. In this case, the measured isotope ratio of a sample is
corrected by simply multiplying it by the α-factor. For multiple
collection, the relative gains of the two collectors must be taken into
account. The gain correction takes an identical mathematical form, and
therefore by referencing it to a known isotope ratio, both gain and IMF
can be corrected at once. Usually, IMF and gain are corrected in a sample
by multiplying the measured, background corrected 18O/16O by the ratio
of the true 18O/16O of a standard to the measured 18O/16O of the standard.
Blocks of data where the IMF changes in step-wise shifts are treated
separately, and corrections are applied when the gain or IMF drifts with
time. Such shifts have proved to be infrequent on SHRIMP II, and are
generally related to definable external factors such as high voltage
fluctuations or temperature variation.
2.9. Precision and accuracy: measurements of MPI-DING reference
material
To assess the instrument performance, including precision,
accuracy, IMF magnitude and sensitivity to the matrix, a series of
analyses was run on six of the MPI-DING natural-silicate-glass
reference materials (Jochum et al., 2006). The MPI-DING glasses
included were the basaltic KL2-G and ML3B-G, komatiitic GOR132-G,
andesitic StHs6/80-G, rhyolitic ATHO-G, and dioritic T1-G. Jochum
et al. (2006) have reported bulk oxygen isotope ratios measured by the
laser fluorination method for each sample to a precision ranging from
0.07 to 0.22‰ (2σ).
Fig. 12. Plot of MPI-DING glass analyses, in chronological order as they were analyzed.
Error bars represent internal precision at 95% confidence. Data is reported in terms of
the IMF as defined in the text.
122
R.B. Ickert et al. / Chemical Geology 257 (2008) 114–128
the dispersion of the mean, as SE. Confidence limits are determined by
multiplying the SD or SE by a Students-t factor for a particular number
of samples. All values of SD and SE here are reported at 95% confidence
limits.
The internal (or “within-spot”) precision is defined as the SE of a
set of isotope ratios measured during the course of a single spot
analysis. For the MPI-DING analyses, the internal precisions ranged
from 0.08 to 0.28‰, with a median value of 0.16‰ (Fig. 13). These
values do not correlate with count rates, secondary ion beam steering
or glass composition. For the count rates and measurement times
during this session, the limit of precision dictated by counting
statistics is approximately 0.04‰, meaning that the internal precisions
measured here approach the theoretically permissible limit (e.g.
Fitzsimons et al., 2000). It is apparent from Figs. 12 and 13 that for all
glasses, the internal precisions are not able to account for the
dispersion in the data. This contrasts with U–Pb analyses of zircon
on SHRIMP II, where internal precisions commonly account for all the
scatter in several measurements of nominally homogeneous reference
zircon.
The spot-to-spot precision is the dispersion of individual analyses
of a “nominally homogeneous” material. It is determined by calculating the SD of the measured reference material oxygen isotope ratios
for a particular session. The designation of material as “nominally”
homogeneous is important, as there is no independent method of
determining oxygen isotope homogeneity on a scale similar to, or
smaller than, the sample size of a SIMS analytical spot. The spot-tospot precisions of the MPI-DING analyses are depicted in Fig. 14. Four
of the glasses have a spot-to-spot precision (i.e., standard deviations at
95% confidence limits) near 0.4‰, one slightly higher at around 0.6‰,
and one much higher at 1.1‰. These values are two to three times
larger than the typical internal precisions (Fig. 13).
There is a large discrepancy between the internal precisions of the
individual analyses and the actual dispersion of data for each glass.
This requires that either 1) the glasses are heterogeneous in oxygen
isotope composition at the scale of 10–100's of µm3, 2) the internal
precisions do not adequately reflect the true uncertainty on the
measurements of the isotopic composition of individual spots, or 3) a
combination of both.
If there is an additional component of analytical uncertainty
associated with the measurement of different spots (case 2, above),
then it should be the case that the dispersion of the data for each glass
should be identical (provided the additional uncertainty is not related
to differing compositions). Bartlett's test for equal variances is a
parametric statistical test that can determine the probability that all of
the samples have equal variances (Snedecor and Cochran, 1989).
When all six samples are taken together, they fail a Bartlett's test
with a probability less than 0.01%, but pass with a probability greater
than 5% when analyses of glass KL2-G are removed. To confirm this
result using another technique that does not require normality, 95%
Fig. 13. Histogram of internal precisions for the MPI-DING glass data. The best precision
possible at these count rates and measurement times is depicted by the light grey band
at ∼0.08‰, and the reproducibility is depicted by a light grey band at ∼ 0.38‰.
Fig. 14. Spot-to-spot precision of MPI-DING glasses. Values plotted are the standard
deviations (at 95% confidence), and the error bars are the 95% confidence limits on these
standard deviations. The error bars have been generated by a bootstrap, and are
asymmetric.
confidence limits on the range in the standard deviations were
constructed using a bootstrap (Wehrens et al., 2000) with 1000
random re-samples. The results of the bootstrap are depicted in Fig. 14
and confirm the results of the Bartlett's test, sample KL2-G clearly has
a significantly different degree of scatter from the remaining six
samples, and the spot-to-spot precisions estimated using those six
samples are statistically equivalent. It is likely that glass KL2-G has
heterogeneous oxygen isotope ratios on a scale and of an amplitude
that can be measured by SIMS.
As has been found previously for SIMS analysis, the internal
precisions significantly underestimate the true analytical uncertainty.
The excess uncertainty is most likely due to slight differences in
charge compensation and secondary ion beam trajectory when the
sample is moved. One method to accommodate the extra uncertainty
would be to expand the internal precisions by an excess scatter term,
for example the square root of the reduced chi-squared (χ2/ν; or
MSWD in the geochronological literature: Wendt and Carl, 1991;
Bevington and Robinson, 2003). This is a useful recourse when the
spot-to-spot precision approaches the internal precision, but for larger
values of the former the excess scatter term dominates. Additionally,
the variation in internal precision (e.g., Fig. 13) probably reflects
simple statistical fluctuation generated by the small number of scans,
and not variation in the confidence with which the mean of the scans
is known. A simpler solution is to use the spot-to-spot precision (the
standard deviation of a series of analyses on a nominally homogeneous standard) as the true uncertainty in the measured isotope
ratio of any given analytical spot. This solution is the one preferred
here, if only for simplicity, and is similar to that in use by other
laboratories (e.g., Page et al., 2007; Trail et al., 2007). We expect that
with improvements to reproducibility our treatment of uncertainties
will evolve.
Finally, the external precision takes into account the uncertainty in
the IMF correction, both in terms of precision of the measured IMF (as
measured on the reference material), and the uncertainty in the true
isotopic composition of the reference material. There are two ways
that this external precision can be presented. If the data of interest are
analytical results from single spots, the external precision on any one
spot is the spot-to-spot precision, the SE of the reference material
analyses, and the uncertainty in the true isotopic composition added
in quadrature. If the data of interest are a pooled set of spots (for
example, during the determination of the oxygen isotopic composition of a homogenous material measured by multiple spots) then the
SE of the sample analyses, the SE of the reference material analyses,
and the uncertainty in the true isotopic composition of the reference
material are added in quadrature. For single analyses, where the true
isotopic composition of the reference material is known to better than
0.1‰, this results in an inflation of the single spot analytical
uncertainty by only 0.02–0.04‰.
R.B. Ickert et al. / Chemical Geology 257 (2008) 114–128
123
2.10. Accuracy and IMF
Table 2
Measured, apparent oxygen isotope ratios in MPI-DING reference glasses
Accurate SIMS isotope ratio measurements depend on how well
the IMF can be measured and how much it differs between the
reference material and material of unknown isotopic composition.
Including the scatter of the individual analyses of each MPI-DING glass
sample, and adding the uncertainty in the true isotopic composition
(Jochum et al., 2006) in quadrature, the precision of the individual
determinations of IMF range from 0.13‰ to 0.30‰, at 95% confidence
limits.
The IMF of the MPI-DING glasses varies strongly with composition.
This illustrates the well-known matrix problem in SIMS, whereby the
IMF varies strongly with the composition of the material that is being
measured (Shimizu and Hart, 1982; Eiler et al., 1997). This relationship
is illustrated in Fig. 15 by plotting the IMF vs. SiO2 concentration. The
IMF varies linearly from rhyolitic to komatiitic compositions (76–
46 wt.% SiO2) over a range of IMF of approximately 7‰.
This relationship emphasizes the need for matrix-matched samples and points the way to a metric for how close in composition
samples and unknowns need to be. The two basaltic composition
glasses, KL2-G (despite probably being slightly heterogeneous) and
ML3B-G provide an example of measuring two matrix-matched
samples. The measured difference in isotope ratios agrees well within
error, suggesting that when using reference material of identical
compositions to unknown material, accuracy can be commensurate
with precision to better than 0.2‰.
Spot Name
Spot Number
18
O−/16O−a 2σb
IMF (‰) 2σb
T1-01
T1-02
T1-03
T1-04
T1-05
T1-06
T1-07
T1-08
T1-09
T1-10
T1-11
T1-12
T1-13
T1-14
Mean
GOR132-01
GOR132-02
GOR132-03
GOR132-04
GOR132-05
GOR132-06
GOR132-07
GOR132-08
GOR132-09
GOR132-10
GOR132-11
GOR132-12
GOR132-13
GOR132-14
Mean
StH-01
StH-02
StH-03
StH-04
StH-05
StH-06
StH-07
StH-08
StH-09
StH-10
StH-11
StH-12
StH-13
StH-14
Mean
ML3B-01
ML3B-02
ML3B-03
ML3B-04
ML3B-05
ML3B-06
ML3B-07
ML3B-08
ML3B-09
ML3B-10
ML3B-11
ML3B-12
ML3B-13
ML3B-14
Mean
ATHO-01
ATHO-02
ATHO-03
ATHO-04
ATHO-05
ATHO-06
ATHO-07
ATHO-08
ATHO-09
ATHO-10
ATHO-11
ATHO-12
ATHO-13
ATHO-14
1
7
13
19
25
31
37
43
49
55
61
67
73
74
0.00203895 ± 15
0.00203886 ± 32
0.00203950 ± 38
0.00203909 ± 32
0.00203902 ± 54
0.00203952 ± 23
0.00203931 ± 45
0.00203997 ± 29
0.00203934 ± 30
0.00203951 ± 45
0.00203984 ± 45
0.00203969 ± 31
0.00203911 ± 33
0.00203990 ± 26
0.00203940 ± 79
0.00204556 ± 23
0.00204587 ± 40
0.00204585 ± 33
0.00204547 ± 22
0.00204605 ± 19
0.00204467 ± 33
0.00204566 ± 33
0.00204582 ± 29
0.00204570 ± 16
0.00204617 ± 24
0.00204568 ± 34
0.00204602 ± 29
0.00204633 ± 39
0.00204590 ± 32
0.00204577 ± 85
0.00203345 ± 33
0.00203347 ± 23
0.00203415 ± 31
0.00203375 ± 25
0.00203418 ± 34
0.00203388 ± 32
0.00203408 ± 32
0.00203390 ± 20
0.00203462 ± 21
0.00203391 ± 38
0.00203350 ± 40
0.00203445 ± 37
0.00203390 ± 50
0.00203385 ± 40
0.00203393 ± 75
0.00204461 ± 56
0.00204513 ± 37
0.00204370 ± 27
0.00204453 ± 28
0.00204432 ± 23
0.00204347 ± 25
0.00204386 ± 32
0.00204466 ± 30
0.00204441 ± 34
0.00204417 ± 35
0.00204535 ± 27
0.00204424 ± 36
0.00204425 ± 37
0.00204375 ± 37
0.00204432 ± 114
0.00202086 ± 30
0.00202151 ± 28
0.00202090 ± 20
0.00202109 ± 29
0.00202096 ± 29
0.00202079 ± 25
0.00202128 ± 28
0.00202134 ± 43
0.00202090 ± 45
0.00202099 ± 28
0.00202115 ± 18
0.00202143 ± 29
0.00202182 ± 25
0.00202195 ± 17
9.23 ± 0.08
9.19 ± 0.16
9.50 ± 0.19
9.30 ± 0.16
9.27 ± 0.27
9.51 ± 0.11
9.41 ± 0.22
9.74 ± 0.14
9.42 ± 0.15
9.51 ± 0.22
9.67 ± 0.23
9.60 ± 0.16
9.31 ± 0.17
9.70 ± 0.13
9.46 ± 0.39
11.51 ± 0.11
11.66 ± 0.20
11.65 ± 0.17
11.47 ± 0.11
11.75 ± 0.10
11.07 ± 0.16
11.56 ± 0.16
11.64 ± 0.14
11.58 ± 0.08
11.81 ± 0.12
11.57 ± 0.17
11.74 ± 0.14
11.89 ± 0.19
11.68 ± 0.16
11.61 ± 0.42
7.92 ± 0.16
7.93 ± 0.12
8.27 ± 0.16
8.07 ± 0.12
8.28 ± 0.17
8.13 ± 0.16
8.23 ± 0.16
8.14 ± 0.10
8.50 ± 0.11
8.15 ± 0.19
7.94 ± 0.20
8.42 ± 0.18
8.14 ± 0.25
8.12 ± 0.20
8.16 ± 0.37
11.21 ± 0.28
11.47 ± 0.18
10.76 ± 0.13
11.17 ± 0.14
11.06 ± 0.12
10.65 ± 0.12
10.84 ± 0.16
11.24 ± 0.15
11.11 ± 0.17
10.99 ± 0.17
11.58 ± 0.14
11.03 ± 0.18
11.03 ± 0.19
10.78 ± 0.19
11.07 ± 0.57
4.60 ± 0.15
4.92 ± 0.14
4.61 ± 0.10
4.71 ± 0.15
4.65 ± 0.14
4.56 ± 0.13
4.81 ± 0.14
4.83 ± 0.21
4.61 ± 0.22
4.66 ± 0.14
4.74 ± 0.09
4.88 ± 0.14
5.07 ± 0.12
5.14 ± 0.08
3. SE Australian granites
The Paleozoic granites of southeastern Australia, where I- and Stype granites were first defined, have been intensely studied for over
three decades (Chappell and White, 1974; Kemp et al., 2007) and have
become textbook examples of granite magmatism (e.g., Winter, 2001).
Despite these studies, however, the nature and origin of these granites
continue to be controversial. Extant problems include the relative
roles of open- and closed-system processes in producing the granites
and their compositional variation, and the genetic relationship
between the I- and S-type granites.
Measurement of oxygen isotope ratios in zircon by ion microprobe
can be a powerful tool in the study of granitic rocks. Although rocks
can easily have their bulk oxygen isotopic compositions modified by
exchange with water during alteration, the chemical stability of zircon
and its very low oxygen diffusion rate (Page et al., 2007) make it highly
resistant to 18O exchange with hydrothermal fluids (King et al., 1997).
In addition, when the slow diffusion of oxygen is coupled with the
long residence time of zircon in magma chambers (Matzel et al.,
2006), it becomes possible to track the isotopic evolution of a
magmatic system through inter- or intra-grain variation in zircon
Fig. 15. Variation of instrumental mass fractionation (IMF) with weight percent SiO2.
The IMF varies linearly with respect to chemical composition. Error bars are at 95%
confidence on the mean, and include the uncertainty in the true isotopic composition of
the sample. Where error bars are not present, they are smaller than the symbols.
2
8
14
20
26
32
38
44
50
56
62
68
75
80
3
9
15
21
27
33
39
45
51
57
63
69
76
81
4
10
16
22
28
34
40
46
52
58
64
70
77
82
5
11
17
23
29
35
41
47
53
59
65
71
78
83
(continued on next page)
124
R.B. Ickert et al. / Chemical Geology 257 (2008) 114–128
Table 2 (continued)
Spot Name
Spot Number
Mean
KL2-01
KL2-02
KL2-03
KL2-04
KL2-05
KL2-06
KL2-07
KL2-08
KL2-09
KL2-10
KL2-11
KL2-12
KL2-13
KL2-14
Mean
6
12
18
24
30
36
42
48
54
60
66
72
79
84
18
O−/16O−a 2σb
IMF (‰) 2σb
0.00202121 ± 78
0.00204377 ± 37
0.00204570 ± 41
0.00204396 ± 45
0.00204438 ± 29
0.00204385 ± 31
0.00204434 ± 27
0.00204501 ± 26
0.00204360 ± 42
0.00204568 ± 25
0.00204632 ± 25
0.00204431 ± 20
0.00204395 ± 22
0.00204560 ± 47
0.00204656 ± 38
0.00204479 ± 217
4.77 ± 0.39
10.52 ± 0.19
11.47 ± 0.20
10.61 ± 0.22
10.82 ± 0.15
10.55 ± 0.16
10.80 ± 0.14
11.13 ± 0.13
10.43 ± 0.21
11.46 ± 0.13
11.77 ± 0.12
10.78 ± 0.10
10.60 ± 0.11
11.42 ± 0.24
11.89 ± 0.24
11.02 ± 1.07
IMF (‰) = ((18O/16Omeasured) / (18O/16Otrue )− 1) ⁎ 1000.
Compositions of MPI-DING glasses reported by Jochum et al. (2006), in δ18O values are:
T1-G= 7.53 ± 0.07; GOR132-G= 8.52± 0.08; StHs6/80-G= 6.12 ± 0.07; ML3B-G= 8.35 ± 0.22;
ATHO-G= 3.20 ± 0.07; KL2-G = 8.63 ± 0.09.
δ 18 O (‰) = (( 18 O/ 16O Sample ) / ( 18O/ 16 OVSMOW )− 1) ⁎ 1000 18 O/ 16OVSMOW = 0.0020052
(Baertschi, 1976).
a
This is the measured 18O/16O, corrected for gain and background but not instrumental
mass fractionation.
b
Internal error.
oxygen isotope ratios. Oxygen isotope ratios are particularly useful for
determining the relative contributions of juvenile and sedimentary
components to granitic magma systems, because the bulk δ18O of the
mantle has a low variance at approximately 5.5‰ (Mattey et al., 1994),
and bulk sedimentary rocks commonly range from 12 to 15‰.
We have measured oxygen isotope ratios in zircon from three
granitoids from the Kosciuszko Batholith in southeastern Australia:
the mafic Blind Gabbro, the intermediate composition I-type
Jindabyne Tonalite, and the S-type Jillamatong Granodiorite. All are
of Siluro–Devonian age (ca. 430–415 Ma) and have been placed in
separate granite suites. The geology and petrography of these samples
are described by Hine et al. (1978) and White and Chappell (1989). The
bulk oxygen isotopic compositions of I- and S-type granites in the
adjacent region have been described by O'Neil and Chappell (1977). Itype granites range from 7.9 to 9.4‰, whereas the S-types have a
narrow range of 9.9 to 10.5‰. These compositions have been
interpreted to reflect the dominantly igneous and sedimentary
precursors of I- and S-type granites, respectively.
Fig. 16. Results of an oxygen analytical session on the ANU SHRIMP II measuring the
isotopic composition of FC1 relative to TEMORA 2. The IMF is very well controlled and
does not drift with time. The data are normalized to a δ18O value of 8.2‰ for TEMORA 2.
The measured difference between FC1 and TEMORA 2 is 2.8 ± 0.3‰ at the 95%
confidence level.
remove the approximately 2 µm deep sputtering pits and implanted
oxygen.
A separate analytical session was run to establish the oxygen
isotopic composition of the FC1 zircon for use as a secondary standard.
TEMORA 2 zircon (δ18O = 8.2‰; Black et al., 2004) was analyzed along
with FC1 zircon (Paces and Miller, 1993). The results are presented in
Table 2 and plotted on Fig. 16. No corrections for IMF/gain drift or EISIE
were necessary. Oxygen isotope analyses of FC1 on SHRIMP II,
normalized to TEMORA 2, yield a mean δ18O value of 5.4 ± 0.3‰.
Laser fluorination oxygen isotope ratio analyses of zircon from similar
rocks by Booth et al. (2005) and Trail et al. (2007) also yield a mean
δ18O value of 5.4‰ (neglecting one low outlier).
3.2. Results
Oxygen isotope ratios were measured in zircon crystals as close as
possible to the spots analyzed for U–Pb, using the techniques
described in Section 2. Each measurement consisted of ten, 10 s
integrations. Baseline count rates on the electrometers were measured at the beginning of the session and have been subtracted from
all the data presented here. Average count rates for 16O and 18O were
2.4·109 cps and 5.2·106 cps, respectively. Corrections for IMF and
detector gains were carried out as described above and all results were
normalized to a δ18O value of 8.2‰ for TEMORA 2. No correction for
IMF drift or gain drift was necessary.
3.1. Analytical procedure
Approximately 2 kg of rock were sampled by sledgehammer or
drill from fresh boulders or roadcuts. Weathered surfaces were
removed, and approximately 0.5 kg of each sample were reduced by
hand to 1 cm sized pieces, followed by fine crushing in a tungsten
carbide swing mill and sieving to b250 µm. Very fine material was
decanted in water and the samples dried under a heat lamp. Zircon
was separated from the remaining powder, using standard heavy
liquid and magnetic techniques. Zircon was hand picked under a
binocular microscope and mounted on double-sided tape prior to
casting in an epoxy “megamount” as described above. Zircon grains
from the granite samples were mounted along with zircon reference
materials TEMORA 2 and FC1, and all grains were contained within an
8 mm by 8 mm square at the centre of the mount.
The mount was coarsely polished to expose the grains just above
their centres and brought to a fine polish using 1 µm diamond paste.
The zircon grains were then all photographed in reflected light,
transmitted light, and by scanning electron microscope cathodoluminescence (SEM-CL). U–Pb ages were measured on SHRIMP II following
the procedures of Williams (1998). The results will be presented in a
separate publication as part of a regional study. The analyzed grains
were then individually imaged by SEM-CL, and then polished again to
Fig. 17. Plot of uncorrected data vs. data that have had the EISIE (Electron-Induced
Secondary Ion Emission) contribution removed. The regression is represented by the
solid grey line and is nearly indistinguishable from a line of unity, indicating an
insignificant effect of the EISIE correction on these analyses.
R.B. Ickert et al. / Chemical Geology 257 (2008) 114–128
Table 3
Measured oxygen isotopic composition of FC1 relative to TEMORA 2
Spot Name
Spot Number
δ18O
FC1z-01.2
FC1z-02.2
FC1z-03.2
FC1z-04.2
FC1z-05.2
FC1z-06.2
FC1z-07.2
FC1z-08.2
FC1z-09.2
FC1z-10.2
TEMz-02.2
TEMz-03.2
TEMz-04.2
TEMz-05.2
TEMz-06.2
TEMz-07.2
TEMz-08.2
TEMz-09.2
TEMz-10.2
TEMz-11.1
1
3
5
7
9
11
13
15
17
19
2
4
6
8
10
12
14
16
18
20
5.73
5.41
5.44
6.03
5.61
5.05
5.19
5.24
5.28
4.98
8.13
7.89
8.54
8.34
7.88
8.13
8.48
7.63
8.41
8.56
2σ (internal)
±
±
±
±
±
±
±
±
±
±
±
±
±
±
±
±
±
±
±
±
0.11
0.23
0.11
0.18
0.14
0.16
0.14
0.11
0.11
0.14
0.16
0.16
0.14
0.11
0.09
0.11
0.16
0.16
0.11
0.11
All data are normalized to a δ18O value of TEMORA 2 of 8.2‰.
δ 18 O (‰) = (( 18 O/ 16O Sample ) / ( 18O/ 16 OVSMOW )− 1) ⁎ 1000 18 O/ 16OVSMOW = 0.0020052
(Baertschi, 1976).
Electron-induced ionization was minimized and kept as constant
as possible during the analyses. On average, the EISIE contributed
0.17% of the total secondary beam, which could require a significant
correction to each spot if peak-stripping were applied, however the
effect was nearly constant for every analysis resulting in very small
relative changes between spots. This can be illustrated simply by
plotting the EISIE corrected and EISIE uncorrected data against each
other (Fig. 17). The slope of this line is 0.99 ± 0.03 with an intercept
of −0.003 ± 0.046, well within error of a slope of unity and an intercept
of zero, demonstrating that there is no systematic difference in
normalized isotopic composition generated by peak-stripping for the
EISIE.
This analytical session included thirteen analyses of TEMORA 2,
and 5 analyses of FC1, both interspersed with analyses of zircon
of unknown isotopic composition (Table 3, Fig. 18). The standard
125
deviation of the TEMORA 2 analyses is 0.34‰ (1σ); therefore the spotto-spot uncertainty for the session (the best estimate of single spot
uncertainty) is 0.75‰ (at 95% confidence). The five analyses of zircon
FC1 yielded a mean δ18O value of 5.3 ± 0.3‰, in excellent agreement
with the value determined above.
Eleven spots were analyzed on zircon from Blind Gabbro sample
MN99-6 (Hine et al., 1978; Table 4). The χ2/ν value for these analyses
is 0.89, indicating that the spot-to-spot uncertainty can account for
analytical scatter in this sample and that 18O/16O is homogenous at the
∼1‰ level. The mean δ18O value of all spots is 5.6 ± 0.3‰ (external
precision at 95% confidence).
Thirteen spots were analyzed on zircon from Jindabyne Tonalite
sample KB22 (Hine et al., 1978). The χ2/ν value is 2.1, suggesting that
the measurements have more scatter than those of the reference
material TEMORA 2. This indicates that the zircon population has a
slightly heterogeneous oxygen isotopic composition just beyond the
1‰ level. The mean δ18O value of all spots is 6.6 ± 0.4‰ (external
precision at 95% confidence).
Eleven spots were analyzed on zircon from the Jillamatong
Granodiorite sample KB165 (Hine et al., 1978). There is considerable
scatter in the analyses: the δ18O values range from 8.2‰ to 10.2‰—a
spread of nearly 2‰—with a mean of 9.1‰. The χ2/ν value for these
analyses is 3.2; clearly indicating that the spot-to-spot uncertainty
does not account for the analytical scatter and that the sample is
heterogeneous.
3.3. Discussion
The gabbro has the most primitive zircon oxygen isotopic
composition, with a δ18O value of 5.6 ± 0.3‰. The whole rock δ18O
value inferred from this and the bulk rock chemical composition
(Lackey et al., 2008) is 6.1‰. A δ18O value of 6.1‰ is outside of the
range of most primitive magmas derived directly from the upper
mantle (Eiler, 2001), but within the range expected for mantle-derived
melts resulting from low degrees of partial melting, or high degrees of
fractional crystallization (Eiler, 2001). One alternative is that the
gabbro may simply be the result of accumulation of minerals from a
more felsic, crustal magma. A second alternative is that a small degree
of crustal contamination of a basaltic magma could be responsible for
the elevated 18O/16O. For example, assimilation of 8% sediment with a
δ18O value of 12‰ by a basaltic magma would raise the δ18O value
Fig. 18. Plot of analyses of δ18O from the I-type Jindabyne Tonalite, S-type Jillamatong Granodiorite, and Blind Gabbro from SE Australia, as well as the standards measured during the
analytical session. Both the tonalite and the grandiorite have less homogeneous oxygen isotopic compositions than the reference materials and the gabbro.
126
R.B. Ickert et al. / Chemical Geology 257 (2008) 114–128
Table 4
Measured oxygen isotope ratios of zircon from granites and reference zircon
Spot Name
Spot Number
δ18O
KB165-07.2B
KB165-08.2B
KB165-09.2B
KB165-10.2
KB165-11.2B
KB165-12.1B
KB165-13.2
KB165-14.2B
KB165-16.1
KB165-17.2
KB165-18.2
KB22-06.1B
KB22-07.2B
KB22-08.2B
KB22-08.3
KB22-09.3B
KB22-10.4
KB22-11.2B
KB22-11.3B
KB22-12.2B
KB22-13.2
KB22-14.1
KB22-15.2
KB22-16.2
MN99-6-02.3
MN-99-6-05.2
MN99-6-07.2B
MN99-6-08.2
MN99-6-09.2B
MN99-6-10.2B
MN99-6-11.2
MN99-6-12.2
MN99-6-13.2
MN99-6-13.2B
MN99-6-15.2
FC1-01.1
FC1-02.1
FC1-02.3
FC1-02.4
FC1-02.5
TEM-01.2
TEM-02.1
TEM-02.2
TEM-03.1
TEM-04.1
TEM-05.1B
TEM-06.1
TEM-06.2
TEM-06.2
TEM-06.3
TEM-06.4
TEM-06.5
TEM-06.6
3
7
11
17
21
26
31
37
41
45
50
4
8
12
13
18
22
27
28
33
38
42
46
51
47
52
5
9
14
19
23
29
39
34
43
1
15
25
32
48
2
6
10
16
20
24
30
35
36
40
44
49
53
8.55
9.47
8.72
8.39
9.82
9.21
10.17
8.18
8.55
9.51
8.96
6.41
5.70
7.06
7.54
6.91
6.71
6.28
6.80
6.41
6.95
6.24
5.78
6.34
5.43
5.51
5.75
5.28
6.06
6.00
4.97
5.57
5.83
5.58
5.21
5.39
5.44
5.54
5.30
5.16
8.52
8.41
7.99
8.43
8.98
8.00
8.14
8.03
7.72
8.28
8.02
8.34
7.75
2σ (internal)
±
±
±
±
±
±
±
±
±
±
±
±
±
±
±
±
±
±
±
±
±
±
±
±
±
±
±
±
±
±
±
±
±
±
±
±
±
±
±
±
±
±
±
±
±
±
±
±
±
±
±
±
±
0.13
0.20
0.35
0.21
0.16
0.33
0.22
0.18
0.12
0.21
0.17
0.21
0.19
0.21
0.12
0.20
0.24
0.17
0.18
0.25
0.23
0.11
0.18
0.15
0.19
0.25
0.21
0.41
0.15
0.13
0.46
0.14
0.20
0.22
0.23
0.25
0.12
0.18
0.17
0.29
0.21
0.26
0.26
0.19
0.22
0.05
0.29
0.11
0.25
0.13
0.13
0.10
0.15
All data are normalized to a δ18O value of TEMORA 2 of 8.2‰.
δ 18 O (‰) = (( 18 O/ 16O Sample ) / ( 18O/ 16 OVSMOW )− 1) ⁎ 1000 18 O/ 16OVSMOW = 0.0020052
(Baertschi, 1976).
from 5.5‰ to 6.1‰. Published initial Nd and Hf isotope data for whole
rocks and zircon from the Blind Gabbro reveal strongly crustal values
of εNdi = −2.5 and εHfi = −2.9 (Kemp et al., 2007), respectively, which is
consistent with our unpublished data of εHfi = −1.5 for zircon from our
sample. These unradiogenic Nd and Hf isotopic compositions are
outside of the normal range for mantle rocks, strongly suggesting the
elevated 18O/16O signature is the result of crustal contamination.
Zircon from the Jindabyne Tonalite has a moderately heavy oxygen
isotopic composition, with a mean δ18O value of 6.6‰. Based on this
value, and the SiO2 concentration in the rock (62 wt.%) the calculated
whole rock δ18O value is 7.9‰. Although no whole rock oxygen isotope
data exist for the Jindabyne Tonalite, other I-type granites in the area
have a range of whole rock δ18O values of 7.9–9.4‰ (O'Neil and
Chappell, 1977), broadly consistent with our calculated value. This is
well in excess of the δ18O expected for mantle-derived rocks, and
requires that the magma that precipitated the zircon includes material
partly or mainly derived from the crust.
Do the oxygen isotope data indicate that the Jindabyne Tonalite
represents a basaltic magma that has assimilated wall rock, or simply a
crustal melt? Assimilation of sedimentary rock by basalt requires the
digestion of a large amount of material, for example, about 37% of
sedimentary rock with a δ18O value of 12‰. Alternatively, partial
melting of meta-basaltic rocks can produce tonalite melts (Rapp and
Watson, 1995) and these may have previously been altered at lowtemperatures which would increase the 18O/16O. High degrees of
assimilation, particularly if they take place in a hot lower crust, are
possible (Thompson et al., 2002), and the nearby presence of the
contaminated Blind Gabbro (if it represents contaminated basaltic
magma) suggests that such processes were indeed ongoing in the
area. In addition, identification of slight isotopic heterogeneity is
strong evidence that some open system processes (such as assimilation or magma mixing) took place in the magma chamber. At this
stage, the evidence suggests that assimilation is the simplest
explanation although compositional data rule out simple bulk mixing
hypotheses (Chappell, 1996).
Zircon from the S-type Jillamatong Granodiorite has the highest
measured 18O/16O, with a range in δ18O values from 8.2 to 10.2‰. The
whole rock δ18O values calculated from these compositions range
from 10.2 to 12.2‰, consistent with the interpretation that the S-type
granites are largely or completely derived from metasedimentary
rocks.
The origin of the isotope disequilibrium in the Jillamatong
Granodiorite, melt precipitated, zircons is intriguing. Some workers
have proposed that S-type granites are generated by interaction
between basaltic and crustal melts (Healy et al., 2004), which could
explain the disequilibrium by zircon precipitation at different stages of
magma mixing or contamination. The local sedimentary rocks have
typical δ18O values of approximately 12‰ (Munksgaard, 1988),
however, requiring that any contribution from a juvenile magma
(δ18O = 5.5‰), be relatively small. When small volumes of hot, mafic
magma encounter large volumes of cool, felsic magma, the mafic
magma is more likely to quench as enclaves than to become
assimilated. Assimilation is therefore unlikely in this case. It is
possible that the zircon have precipitated from an inhomogeneous
magma, or a series of closely related magmas that aggregated at the
site of magma emplacement. This is likely to happen in crustal melting
situations where the source regions are heterogeneous, and the
magmas are relatively cool, crystal rich, and viscous (e.g., Bindeman
et al., 2008).
A large diversity of primary magma δ18O values, from a range of
rock types, has been identified by analysis of oxygen isotope ratios in
zircon using the ANU SHRIMP II. The high spatial resolution of the ion
probe can be used to identify pristine regions of crystals that may best
preserve information on the original magmatic δ18O values, and can
be utilized to detect small amounts of crustal contamination. This was
the case for the altered Blind Gabbro, where the zircon δ18O value
preserves the original magmatic oxygen isotope ratios and the slight
elevation of 18O/16O points to a small degree of crustal contamination.
The contamination must have occurred prior to the crystallization of
zircon. The ability to identify intercrystal isotope disequilibrium, such
as in the case of the Jindabyne Tonalite, can be used to resolve
competing petrologic hypotheses such as closed vs. open system
behaviour of magma chambers. In this case, a combination of elevated
δ18O values above the mantle mean, and slight heterogeneity, suggest
that open system processes played a role in the tonalite's history. The
oxygen isotope history of zircon in the S-type granite is more complex,
although individual δ18O values for zircon clearly indicate a sedimentary source. The origin of the intercrystal isotope variation is unclear. It
may simply be vestigial from a similarly heterogeneous source.
R.B. Ickert et al. / Chemical Geology 257 (2008) 114–128
4. Conclusions
Technical improvements to the SHRIMP II ion microprobe have
made it possible to use the instrument to measure the oxygen isotopic
composition of insulating materials on a 25 µm scale with better than
0.4‰ precision and accuracy. Key aspects of this achievement have
included the introduction of a Cs+ source, a multiple collector, charge
neutralization using an oblique-incidence electron gun, Helmholtz coils,
and the development of new sample mounts that reduce geometry
effects. New oxygen isotope ratio measurements from zircon in igneous
rocks from southeastern Australia shed new light on an old debate,
hinting at a complex history for the I-type granite magmas and
confirming the heterogeneous sedimentary origin of S-type magmas.
Acknowledgements
Funding from the Australian Research Council (including DP0559604)
supported the experiments leading to the technical advances reported in
this paper. Hardware modifications were funded by the Australian
National University. RBI was supported by an Endeavour International
Postgraduate Research Scholarship awarded by the ANU and a
postgraduate scholarship provided by the Natural Sciences and Engineering Research Council of Canada. Marc Norman is thanked for
providing the sample of Blind Gabbro, and B. Stoll and K. P. Jochum kindly
provided the samples of MPI-DING glass. Bruce Chappell is thanked for
extensive discussions and comments on an earlier version of the
manuscript. Harold Persing of Australian Scientific Instruments is
thanked for his efforts in helping to develop the hardware. Richard
Hervig and John Valley are thanked for constructive reviews that
substantially improved the clarity of the paper, and Roberta Rudnick is
thanked for patient editorial handling.
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