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Geochimica et Cosmochimica Acta 77 (2012) 415–431
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26
Al–26Mg deficit dating ultramafic meteorites and silicate
planetesimal differentiation in the early Solar System?
Joel A. Baker a,⇑, Martin Schiller a,b, Martin Bizzarro b
a
School of Geography, Environment and Earth Sciences, Victoria University of Wellington, P.O. Box 600, Wellington 6014, New Zealand
b
Centre for Star and Planet Formation, Natural History Museum of Denmark, University of Copenhagen, Øster Voldgade 5-7,
Copenhagen DK-1350, Denmark
Received 10 June 2011; accepted in revised form 17 October 2011; available online 21 October 2011
Abstract
Meteorites with significantly sub-chondritic Al/Mg that formed in the first 2 million years of the Solar System should be
characterised by deficits in the abundance of 26Mg (d26Mg*) due to the absence of in-growth of 26Mg from the decay of shortlived 26Al (t1/2 = 0.73 Myr). However, these 26Mg deficits will be small (d26Mg* >0.037&) even for material that formed at
the same time as the Solar System’s oldest solids – calcium–aluminium-rich inclusions – and thus measurement of these deficits is analytically challenging.
Here, we report on a search for 26Mg deficits in three types of ultramafic meteorites (pallasites, ureilites and aubrites) by
multiple-collector inductively coupled plasma mass spectrometry. A range of analytical tests were carried out including analysis of: (1) a range of synthetic Mg solution standards; (2) Mg gravimetrically doped with a high purity 26Mg spike; (3) Mg
cuts collected sequentially from cation exchange separation columns with fractionated stable Mg isotope compositions; (4)
Mg separated from samples that was bracketed by analyses of both DSM-3 and Mg separated from a natural olivine sample
subjected to the same chemical processing as the samples. These tests confirm it is possible to resolve differences in d26Mg*
from the terrestrial materials that are 60.005&. However, if Mg yields from chemical separation are low or an inappropriate
equilibrium-isotopically fractionated standard is used this will generate analytical artefacts on d26Mg* when this is calculated
with the kinetic/exponential mass fractionation law as is the case when correcting for instrumental mass bias during mass
spectrometric analysis.
Olivine from four different main group pallasites and four bulk ureilites have small deficits in the abundance of 26Mg
with d26 MgDSM-3 ¼ 0:0120 0:0018& and d26 MgDSM-3 ¼ 0:0062 0:0023&, respectively, relative to terrestrial olivine
(d26 MgDSM-3 ¼ þ0:0029 0:0028&). Six aubrites have d26 MgDSM-3 ¼ þ0:0015 0:0020&, which is identical to terrestrial
olivine.
Model ages from these deficits can be calculated by assuming that 26Al was homogeneously distributed in the planetesimalforming regions of the proto-planetary disc at the same level as calcium–aluminium-rich inclusions (CAIs). The absence of
26
Mg deficits in aubrites, means these can only be constrained to have formed relatively late >2.9 Myr after CAI formation.
Model ages calculated from pallasite olivine deficits would suggest that pallasite olivine crystallised and was diffusively isolated on its parent body 1:24þ0:40
0:28 Myr after formation of CAIs. Similarly, ureilites would have experienced silicate partial
melting and lowering of their bulk Al/Mg ratios 1:9þ2:2
0:7 Myr after CAI formation. The model ages for silicate differentiation
on the main group pallasite parent body are intermediate between those for metal-silicate fractionation for core formation
obtained from magmatic iron meteorites and those for asteroidal silicate magmatism obtained from basaltic meteorites.
Ó 2011 Elsevier Ltd. All rights reserved.
⇑ Corresponding author. Tel.: +64 4 463 5493; fax: +64 4 463 5186.
E-mail address: [email protected] (J.A. Baker).
0016-7037/$ - see front matter Ó 2011 Elsevier Ltd. All rights reserved.
doi:10.1016/j.gca.2011.10.030
416
J.A. Baker et al. / Geochimica et Cosmochimica Acta 77 (2012) 415–431
1. INTRODUCTION
A variety of long- (absolute) and short-lived (relative)
chronometers are used to date meteorites and their components as a result of the processes of solid formation and
planetary accretion and differentiation in the early Solar
System. In particular, over the past several decades, application of absolute Pb–Pb and relative 53Mn–53Cr,
182
Hf–182W and 26Al–26Mg dating techniques have led to
an increasingly clearer picture of these timescales (e.g.,
Lugmair and Galer, 1992; Lugmair and Shukolyukov,
1998; Srinivasan et al., 1999; Amelin et al., 2002; Kleine
et al., 2002; Yin et al., 2002; Bizzarro et al., 2005; SpivakBirndorf et al., 2009; Wadhwa et al., 2009). However, a
number of types of meteorites and their components are
not easily dated with conventional application of these isotopic systems. Examples of meteorites that are difficult to
date include those where: (1) the meteoritic material may
contain very low lithophile trace element abundances characterised by a low ratio of the parent (P) radioactive isotope
to the daughter (D) radiogenic isotope. (2) Particularly in
combination with (1) some isotopic systems are highly sensitive to the affects of terrestrial contamination, such as the
Pb–Pb chronometer (Torigoye-Kita et al., 1995a; Amelin,
2006). (3) Co-existing phases in a meteorite with high and
low P/D ratios have re-equilibrated due to slow cooling
of the meteorite parent body and/or due to later thermal
or shock events, especially where the high P/D phase has
high diffusivity for the element of the daughter isotope.
Numerous examples of meteorite dating studies with the
53
Mn–53Cr, 182Hf–182W and 26Al–26Mg chronometers have
interpreted the obtained ages to reflect secondary events
rather than the primary crystallization age of a meteorite
(or its components) (e.g., Wadhwa et al., 2003; Shukolyukov and Lugmair, 2004; Kleine et al., 2005a; Touboul
et al., 2009).
26
Al–26Mg dating has been used as a relative dating tool
for meteorites since the first demonstration that live 26Al
was present in the Solar System’s oldest dated solids (calcium–aluminium-rich inclusions; CAIs) when they formed
(Lee et al., 1977; Gray and Compston, 2004). Since then
both bulk analytical methods applied to whole rock samples and mineral separates (thermal ionisation and plasma
source mass spectrometry) and in situ (secondary ionisation
and laser ablation plasma source mass spectrometry) analytical methods have utilised the 26Al–26Mg chronometer
to date solid formation (CAIs and chondrules) and planetesimal magmatism (e.g., Russell et al., 1996; Kita et al.,
2000; Bizzarro et al., 2005; Simon et al., 2005; Young
et al., 2005; Jacobsen et al., 2008; Spivak-Birndorf et al.,
2009). In these cases, the 26Al–26Mg chronometer has been
primarily used to date meteorites or their components that
have high P/D (i.e., Al/Mg) ratios.
The initial abundance of 26Al in the Solar System is sufficiently high that meteorites which formed very early in the
Solar System with markedly sub-chondritic Al/Mg (i.e.,
27
Al/24Mg 0.101; Lodders, 2003) might have less radiogenic 26Mg as compared to the bulk Solar System
(Fig. 1). Fig. 1 shows a Mg isotopic evolution curve for
Fig. 1. d26Mg* isotopic evolution of the Solar System in the first
5 million years after CAI formation. Meteorites or meteoritic
material characterised by sub-chondritic Al/Mg (27Al/24Mg
<0.101) will be characterised by resolvable d26Mg* deficits if they
formed within 2 million years of CAIs. The isotopic evolution
curve shown is calculated using the initial 26Al value for the Solar
System and its proto-planetary disc based on the assumption that
the planetesimal-forming region of the proto-planetary disc had the
same initial 26Al/27Al as indicated by a regression through Mg
isotope data for chondrites and CAIs (Jacobsen et al., 2008;
Schiller et al., 2010a).
the Solar System based on an initial 26Al/27Al value of
5.21 105 derived from a regression through high precision Mg isotope data for CAIs and bulk chondrites (Jacobsen et al., 2008; Schiller et al., 2010a) and the chondritic Al/
Mg ratio, which shows it is possible to hypothesize that the
early Solar System evolved from an initial d26Mg* deficit of
0.037& to essentially its present value within 5 million
years.
While these potential d26Mg* deficits are small, improved analytical methods, and evidence that planetesimals
accreted and differentiated very early in the Solar System
(e.g., Kleine et al., 2005b) make it plausible that it will be
possible to detect such deficits in appropriate meteoritic
material. This would provide a potential dating methodology for high-Mg meteoritic material in a manner that is
analogous to 182Hf–182W dating of metal phases in meteorites with Hf/W 0 (e.g., Kleine et al., 2005b; Markowski
et al., 2006a,b; Schersten et al., 2006). This type of approach has recently been utilised by Villeneuve et al.
(2010) to date chondrule formation, and Mg-rich olivines
in chondrites and the Eagle Station pallasite by ion probe
methods.
Herein we describe the methods and results of a search
for deficits in the abundance of 26Mg in three classes of
meteorites (pallasites, ureilites and aubrites) with subchondritic Al/Mg ratios. We show that it is possible to
resolve small variations in d26Mg* in meteorites compared to terrestrial material and chondrites and potentially place some new age constraints on the formation
of pallasites, ureilites and aubrites, and silicate planetary
differentiation of their parent bodies in the early Solar
System.
26
Al–26Mg deficit dating ultramafic meteorites
2. STANDARDS, SAMPLES AND ANALYTICAL
TECHNIQUES
2.1. Standards and samples
A range of synthetic Mg solution standards and Mg separated from an in-house olivine standard (J11) taken from a
spinel peridotite were analysed in this study. The primary
Mg solution standard used for bracketing analyses of other
Mg solution standards as well as Mg separated from terrestrial olivine and meteorite samples was DSM-3 (Galy et al.,
2003). Various ICP-MS Mg solution standards (Aristar,
Alfa Aesar 1, and Alfa Aesar 2) and SRM980 were analysed versus DSM-3. Aristar Mg gravimetrically spiked
with a high purity (>99.5%) 26Mg spike and Aristar Mg
subjected to various chemical separation techniques was
also analysed versus unspiked/unprocessed Aristar Mg
(Section 2.5).
Mg separated from olivine minerals in samples J11, JK3
and JB281 was analysed versus the DSM-3 Mg standard
and, in a number of cases, Mg separated from J11 mantle
olivine was also used as the bracketing standard for Mg isotope analysis of Mg separated from these terrestrial olivines
and the meteorite samples. Where Mg separated from J11
mantle olivine was used as the bracketing standard, Mg
was separated from J11 olivine in exactly the same fashion
as the terrestrial olivines or meteorite samples being analysed.
J11 is an anhydrous spinel peridotite collected from a
Plio-Quaternary intraplate volcanic cone located in Jordan
(Shaw et al., 2007), whereas JK3 is a hydrous amphibolebearing spinel peridotite collected from Plio-Quaternary
intraplate volcanism located in Ataq (southern Yemen)
(Baker et al., 1998) and JB281 are olivine phenocrysts from
a near-primary continental flood basalt sourced from the
Afar mantle plume erupted in Yemen during the Oligocene
(Baker et al., 1996).
Three different types of meteorites with sub-chondritic
Al/Mg were analysed in this study – olivine from main
group pallasites, and bulk ureilite and aubrite samples.
Main group pallasites are stony-iron meteorites containing
large mm-sized olivine crystals set in evolved iron–nickel
metal that are thought to represent the core-mantle boundary of differentiated planetesimals (Wasson and Choi,
2003). Ureilites are coarse-grained meteorites composed
primarily of olivine and pyroxene, but also containing
Fe–Ni metal, sulphide phases and various forms of carbon
(e.g., Goodrich, 1992). Ureilites are highly depleted in
incompatible lithophile elements and variably depleted in
siderophile elements (e.g., Boynton et al., 1976; Goodrich,
1992; Mittlefehldt, 2007, reference therein). While ureilites
exhibit some features typical of both primitive and differentiated meteorites and their precise origin remains enigmatic,
they are generally considered to represent partial (s)melting
residues of a differentiated planetesimal or planetesimals
(Warren et al., 2006). Aubrites are coarse-grained meteorites primarily composed of variably brecciated Fe-free orthopyroxene with associated small and varying amounts of
plagioclase, high-Ca pyroxene and forsterite as well as a
suite of accessory metal and sulphide phases (Mittlefehldt,
2007, reference therein). Aubrites are highly depleted in
417
both lithophile and siderophile trace elements and,
although their origins are also unclear, are interpreted as
being coarse-grained igneous cumulates from a highly reduced, differentiated planetesimal.
Magnesium separated from olivine from four main
group pallasites (Admire, Brenham, Esquel and Molong)
was analysed in this study. These pallasites contain olivine
with a restricted range of Fo contents (87–89; Wasson and
Choi, 2003). Small sub-mm-sized fragments of olivine devoid of chromite inclusions were hand-picked from the pallasite meteorites (and also terrestrial mantle and basalt
samples) under a binocular microscope. Prior to digestion,
the olivine samples were washed with ultra-clean water
(>18.2 MX) and gently acid washed with cold 2 M HCl
for 5–10 min to remove any secondary material resulting
from oxidation of iron–nickel metal and/or the olivine.
Bulk samples of four ureilites were analysed in this
study. SAH98505 is a coarse-grained ureilite dominated
by olivine (Fo81) and pigeonite (Fs12.9) (Grossman, 1999).
El Gouanem is a ureilite find from Morocco and has a typical ureilitic texture and is primarily composed of olivine
with Fo76–78 (Grossman and Zipfel, 2001). NWA2234 is a
crystalline ureilite composed of coarse, shocked dusty olivine (cores Fo82–92, rims Fo94–98) and pigeonite (Fs17) (Russell et al., 2004). NWA766 is a ureilite containing 80%
olivine (Fo76) and 20% pigeonitic pyroxene (Fs18.7), but
marked by the presence of a high-Si–Al glass (Skirdji and
Warren, 2001).
Bulk samples of six aubrites were analysed – Norton
County, Pena Blanca Spring, Mt. Egerton, Shallowater,
Cumberland Falls and Bishopville. All of these aubrites
are dominated by unbrecciated (Mt. Egerton and Shallowater) to variably brecciated Fe-free enstatite pyroxene
crystals, although Cumberland Falls contains some
unequilibrated chondritic material (Neal and Lipschutz,
1981) and Bishopville contains a much higher modal abundance of plagioclase than the other aubrites (Watters and
Prinz, 1979).
Representative fragments of ca. 100–200 mg of the ureilites and aubrites were hand-picked under a binocular
microscope. These fragments were then crushed to a powder with an agate mortar and pestle.
2.2. Sample digestion and chemical separation of Mg
Samples were digested in a 3:1 mixture of concentrated
HF and HNO3 acid in savillex Teflon capsules on a hotplate at 130 °C. Approximately 5–10 mg of olivine from
the terrestrial samples and pallasites was digested, whereas
ca. 50 mg aliquots of powdered ureilite or aubrite material
was digested. After evaporation of the HF–HNO3 acid,
samples were sequentially refluxed and evaporated with
concentrated HNO3, 7 M HCl and aqua regia to bring them
fully into solution. Samples were finally converted to chloride form by evaporation of 7 M HCl prior to dissolution in
concentrated HCl for the first Mg chemical separation step.
All acids used in this study were high purity Seastar acids,
where necessary, diluted with >18.2 MX ultra-clean water.
Chemical separation techniques for purification of Mg
utilised in this study have been previously described in
418
J.A. Baker et al. / Geochimica et Cosmochimica Acta 77 (2012) 415–431
detail by Handler et al. (2009) and Schiller et al. (2010a,b)
and are only briefly described here. An amount of sample
equivalent to ca. 1 mg of Mg was subjected to up to five
chemical separation steps comprising: (1) anion exchange
separation (0.5 mL Bio-Rad AG1-X4 resin) of Fe in concentrated HCl whereby Fe is retained on the resin; (2) separation of Ca in 3 M HNO3 on 0.25 mL Eichrom DGA
resin whereby Ca is retained on the resin; (3) cation exchange separation of other major and trace elements with
the exception of Mn and Ni in 1 M HNO3–0.1 M HF on
1 mL of Bio-Rad AG50W-X8 200–400 mesh resin; (4) cation exchange separation of Mn in 0.5 M HCl–95% acetone
on 1 mL of Bio-Rad AG50W-X8 200–400 mesh resin and
(5) separation of Ni using 1 mL of Eichrom Ni-spec resin.
Different types of samples were subjected to different
chemical separation steps. The aubrites and relatively Nirich ureilites were subjected to the most extensive chemical
separation steps (aubrites – steps 1–4 above; ureilites steps
1–5 above) to ensure the complete removal of matrix elements. However, the terrestrial and pallasite olivines have
relatively simple compositions compared to the aubrites
and ureilites. Thus, the olivines were analysed after just step
(1), steps (1–3) and steps (1–4) to assess the extent to which
different chemical separation procedures affected the Mg
isotope results. When Mg separated from J11 was used as
a bracketing standard, it was treated to exactly the same
Mg chemical separation procedure as the same being analysed to minimise any potential analytical artefacts.
All Mg separates were screened for the presence of matrix elements prior to Mg isotope analysis. Mg extracted
from olivines after the first anion exchange step was
>98.5% (terrestrial olivine) and >99.0% (pallasite olivine)
pure. Mn and Ni were the most abundant matrix elements
remaining (1.0–1.5%) after this first chemical separation
step, along with very minor amounts of Al, Ca and Cr
(<0.1% total), although no Ni was present in the pallasite
olivine Mg due to the very low Ni contents of pallasite olivine. After complete processing, Mg purity was always
>99.5% in all the analysed terrestrial and meteoritic samples, with only small amounts of Ni (<0.5%) remaining in
the Mg separated from the terrestrial olivines. Mg procedural blanks were always <0.001% of the processed Mg
for each sample. All of the chemical separation steps were
checked to ensure that they resulted in 100% Mg yields.
2.3. Al/Mg ratio measurements by ICP-MS
Al/Mg ratios were measured on aliquots of dissolved
samples taken before chemical separation of Mg with an
Agilent 7500CS ICP-MS in Victoria University of Wellington’s Geochemistry Laboratory using He in a collision cell
to minimise interferences on Al and Mg isotopes (27Al,
24
Mg and 25Mg). Sample analyses were bracketed with
analyses of a gravimetrically prepared Al/Mg = 0.1 solution made from Aristar single element ICP-MS solutions.
The error assigned to the Al/Mg ratio is ±2% (2 sd) based
on repeated measurements of USGS basaltic rock standards BCR-2 and BHVO-2 (Schiller et al., 2010b). Al/Mg
ratios of the pallasite olivines were not measured precisely
as these were always <0.001.
2.4. Mg isotope analyses by MC-ICP-MS
Mg isotope ratios were measured in pseudo-high-resolution mode with a Nu Plasma MC-ICP-MS in Victoria University of Wellington’s Geochemistry Laboratory. Mg
solutions containing ca. 1–3 ppm Mg were introduced into
the plasma with a DSN-100 desolvating nebuliser system.
The mass spectrometer was operated at a resolution of ca.
2000–2500, which enables resolution of polyatomic interferences on the high mass side (e.g., 12C2+, 12C14N+) of Mg.
24
Mg, 25Mg and 26Mg were either monitored by the L5,
Ax and H6 Faraday collectors equipped with 1011 X resistors or by L4, Ax and H5 collectors where the L4 collector
was equipped with a 1010 X resistor allowing larger ion
beams of ca. 25–50, 3–6 and 3–6 V to be measured on
masses 24, 25 and 26 respectively. Results using the latter
collector configuration do not differ significantly from the
former configuration except that, in some analytical sessions, internal errors were improved by about a factor of
1.5 when measuring the larger ion beams.
A single Mg isotopic analysis comprises a total of 480 s
of baseline measurements and 1600 s of data acquisition in
four blocks (four blocks of 80 5 s integrations). Sample
analyses were either bracketed by analyses of the DSM-3
standard or by Mg separated from the J11 in-house olivine
standard. J11 olivine has a stable Mg isotopic composition
that is slightly lighter than DSM-3 and more representative
of terrestrial Mg than DSM-3 (Handler et al., 2009). We
used Mg separated from J11 as the bracketing standard
for some of our analyses for two reasons. Firstly, it is possible (but not known) that the slight isotopic difference between DSM-3 and Earth could be the result of equilibrium
isotopic fractionation processes (Young and Galy, 2004). If
so, pure equilibrium fractionation would produce very marginally erroneous mass-bias-corrected 26Mg* values for the
mass-independent abundance of 26Mg when calculated
using the kinetic (=exponential) fractionation law by
0.004& per 0.1& difference in the stable isotopic difference
(d25Mg) between the sample and bracketing standard. Secondly, we used Mg separated from J11 as the bracketing
standard for some analyses as this meant that the sample
and standard had both experienced exactly the chemical
separation procedures, providing an additional test of our
analytical methodology.
The mass-independent abundance of 26Mg (d26Mg*) was
calculated by internally normalising the 26Mg/24Mg to
25
Mg/24Mg = 0.12663 (Catanzaro et al., 1966) using the
exponential mass fractionation law (b = 0.511) and calculating the difference between this value for the sample and
the average value of the bracketing standards in the per
mil notation (&). Stable Mg isotope data (d25Mg) are reported in the per mil notation as the difference between
the sample and the average value of the bracketing
standards.
Single Mg isotope analyses have uncertainties (2 se) on
d26Mg* that are ±0.021& to ±0.012& when the uncertainties on the bracketing standards are quadratically incorporated into the error on the sample. Each Mg isotope
analysis presented in Tables 1–4 represent the weighted
mean of 2–38 such measurements resulting in final 2 se
Table 1
Mg isotope data for standard solutions (Aristar, Alfa Aesar, SRM980) and terrestrial olivine separated from mantle (J11, JK3) and basalt (JB281) samples.
Solution standards
Average & 2 se
Weighted mean & 2 se
0.0178
0.0168
0.0149
0.0014
0.0104
0.0166
0.0305
0.1928
0.0005
0.0065
0.0026
0.0063
0.0040
0.0060
0.0076
0.0041
0.0044
0.0029
0.0022
0.0051
0.0049
0.0037
0.0171
0.0169
0.0154
0.0027
0.0107
0.0171
0.0310
0.1930
0.0003
0.0069
0.0019
0.0015
0.0007
0.0011
0.0063
0.0030
0.0016
0.0001
0.0004
0.0038
0.0069
0.0033
0.0042
0.0042
0.0081
0.0048
0.0047
0.0021
0.0007
0.0025
0.0009
Aristar
Alfa Aesar 1
Alfa Aesar 2
SRM980 (NZ)
Aristar (d26Mg* = 0.0100&)A
Aristar (d26Mg* = 0.0200&)
Aristar (d26Mg* = 0.0300&)
Aristar (d26Mg* = 0.2000&)
Aristar (column processed)B
Aristar (column processed)
Aristar (column processed)
Terrestrial mantle and basalt olivineC
J11 digestion 1
J11 digestion 2
J11 digestion 2
J11 digestion 3
J11 digestion 4
J11 digestion 5
JK3
JB281
J11 olivine mean versus DSM-3
J11/JB281/JK3 olivine mean versus J11
±2 se (&)
d26Mg* (&)
d25Mg (&)
±2 se (&)
d26Mg (&)
±2 se (&)
n
Bracketing standard
ChemistryD
0.0055
0.0093
0.0083
0.0072
0.0060
0.0037
0.0080
0.0100
0.0071
0.0072
0.0057
0.819
0.405
1.716
2.329
0.039
0.125
0.114
0.079
0.032
0.024
0.032
0.029
0.117
0.071
0.072
0.081
0.065
0.182
0.055
0.131
0.066
0.036
1.584
0.781
3.334
4.555
0.066
0.263
0.251
0.039
0.071
0.056
0.067
0.058
0.229
0.137
0.138
0.153
0.127
0.354
0.105
0.241
0.134
0.070
7
4
5
7
8
22
5
3
6
7
10
J11 olivine
J11 olivine
J11 olivine
J11 olivine
Aristar
Aristar
Aristar
Aristar
Aristar
Aristar
Aristar
None
None
None
None
None
None
None
None
a
a, dga, c
a, dga, c
0.0021
0.0001
0.0013
0.0077
0.0028
0.0016
0.0006
0.0007
0.0050
0.0072
0.0086
0.0057
0.0068
0.0062
0.0059
0.0058
0.103
0.106
0.195
0.065
0.030
0.063
0.033
0.056
0.075
0.068
0.041
0.026
0.055
0.251
0.043
0.051
0.191
0.209
0.389
0.123
0.059
0.125
0.061
0.116
0.153
0.135
0.079
0.051
0.113
0.486
0.079
0.105
22
6
5
11
7
10
9
8
DSM-3
DSM-3
DSM-3
DSM-3
DSM-3
J11 olivine
J11 olivine
J11 olivine
a
a
a,
a,
a,
a,
a,
a,
0.0029
0.0005
0.0028
0.0034
0.100
0.029
0.055
0.062
0.194
0.060
0.111
0.121
±2 se (&)
dga,
dga,
dga,
dga,
dga,
dga,
c
c, cMn
c, cMn
c, cMn
c
c
Al–26Mg deficit dating ultramafic meteorites
d26Mg* (&)
26
Sample
A
Aristar solutions with gravimetrically prepared artificial d26Mg* excesses.
About 1000 lg of Aristar Mg standard solution chemically processed through the Mg chemical separation procedure listed.
C
J11 = mantle olivine from an anhydrous spinel peridotite (Jordan); JK3 = mantle olivine from a hydrous spinel peridotite (Yemen); JB281 = olivine phenocrysts from an Oligocene continental
flood basalt (Yemen).
D
a = Anion chemical separation (Fe); dga = TODGA Eichrom separation (Ca); c = cation exchange separation (most elements except Mn and Ni); cMn = cation exchange separation in HCl/
acetone (Mn).
B
419
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J.A. Baker et al. / Geochimica et Cosmochimica Acta 77 (2012) 415–431
Table 2
Mg isotope data for Aristar Mg standard solution isotopically fractionated on cation exchange columns.
Sample
d26Mg* (&)
d25Mg (&)
±2 se (&)
d26Mg (&)
±2 se (&)
Aristar (column cut; 50–70 mL)A
Aristar (column cut; 70–90 mL)
Aristar (column cut; 90–110 mL)
Weighted mean & 2 se
0.0270
0.0130
0.0060
0.0100
0.0270
0.0120
1.223
0.164
0.679
0.060
0.152
0.086
2.359
0.309
1.302
0.108
0.295
0.163
2
3
3
38 mLB
43 mL
49 mL
55 mL
66 mL
88 mL
0.0315
0.0160
0.0040
0.0129
0.0210
0.0050
1.383
0.447
0.138
0.265
0.763
0.188
0.076
0.035
0.051
0.048
0.059
0.042
2.680
0.859
0.282
0.506
1.470
0.359
0.150
0.067
0.093
0.095
0.110
0.084
14
15
5
10
9
10
±2 se (&)
0.0055
0.0055
0.0130
0.0061
0.0066
0.0062
n
A
About 20 mL aliquots of Aristar Mg sequentially collected from 5000 lg of Aristar Mg standard passed through a cation exchange column
with a resin bed of 4.5 mL in 1 M HNO3.
B
The 1–2 mL aliquots of Aristar Mg sequentially collected from 10,000 lg of Aristar Mg standard passed through a cation exchange
column with a AG50W-X8 resin bed of 4.5 mL in 1 M HCl.
analytical uncertainties of ±0.013& to ±0.004&. While the
errors for the data presented here are internal errors, the
external reproducibility on d26Mg* was typically found to
be 50% greater than internal errors based on repeated measurements of samples, terrestrial standards and standards
with gravimetrically prepared excesses in 26Mg (herein;
Schiller et al., 2010b). Thus we estimate that the external
reproducibility of sample analyses is a factor of 1.5 greater
than internal errors.
A typical analysis of a meteorite sample comprises the
mean of 6–12 individual measurements carried out over
an 8–14 h period consuming about 30–90 lg of Mg. In a
number of cases, this type of analysis would then be repeated on either remaining Mg left over from the first 6–
12 analyses on another day in another analytical session,
or on Mg separated in a new chemical separation chemistry
from remaining digested material of the sample, or of Mg
separated from the same sample in a new chemical separation chemistry where new material was digested a second or
third time. In the data presented in Tables 1–4, n refers to
the number of repeat measurements carried out on each
solution. When a sample was analysed in a number of different ways and results pooled into a final mean, n is expressed as, for example 35/4/3 (Table 1; Admire pallasite
olivine measured versus DSM-3). In this case, three different separates of olivine from the Admire pallasite were digested, subjected to chemical separation on Mg three
occasions utilising three different types of chemical separation, and analysed in total 35 times in four analytical
sessions.
2.5. Analytical tests
Six analytical tests were carried out on standards, terrestrial samples and meteorite samples to assess potential analytical artefacts and demonstrate the precision and accuracy
of the presented Mg isotope data: (1) a range of synthetic
Mg solution standards were analysed versus DSM-3. (2)
Aristar Mg doped with a high purity 26Mg spike to produce
26
Mg* anomalies in the range of 0.010–0.200& with an
accuracy of <5% was analysed against undoped Aristar
Mg. (3) Aristar Mg subjected to various parts of the Mg
chemical separation procedures was analysed versus unprocessed Aristar Mg. (4) Aristar Mg cuts collected sequentially from cation exchange separation columns with
fractionated stable Mg isotope compositions were analysed
to examine the effects of incomplete Mg recovery on measurements of the mass-independent abundance of 26Mg.
In this experiment, a larger volume of resin (4.5 mL) was
utilised than in the processing of samples, in order to ensure
that isotopically fractionated cuts of Mg would be obtained. (5) Mg separated from three different terrestrial olivine samples and some meteorite samples was analysed
versus both DSM-3 and Mg separated from the in-house
J11 mantle olivine standard. (6) The Mg isotopic composition of both the mantle olivine (J11) and pallasite olivine
samples were measured after different stages (anion exchange separation of Fe ± TODGA separation of Ca ±
cation exchange separation of most major and trace
elements ± cation exchange separation of Mn in 0.5 M
HCl–95% acetone) of chemical separation of Mg to assess
the extent to which the progressive removal of matrix elements affected the results.
3. RESULTS
3.1. Analytical tests and analyses of solution and mineral
standards
3.1.1. Analysis of Mg solution standards versus DSM-3
Four Mg solution standards were analysed versus DSM3 and yielded light stable Mg isotopic compositions of
d25Mg = 0.41& to 2.33& (Table 1). With the exception
of the most fractionated standard (SRM980) all the other
Mg solution standards show small apparent excesses in
the abundance of 26Mg with respect to DSM-3 of
d26Mg* = +0.0154& to +0.0171& (Table 1 and Fig. 2).
3.1.2. Analysis of Aristar Mg doped with a high purity 26Mg
spike
Four Aristar Mg standard solutions were gravimetrically doped with >99.5% pure 26Mg to create solutions with
Table 3
Mg isotope data for olivine from pallasite meteorites.
Sample
Type
27
Al/24Mg
Fo87.9
“
“
“
“
“
“
“
0.000
“
“
“
“
“
“
“
Pallasites
Admire
Admire
Admire
Admire
Admire
Admire
Admire
Admire
digestion
digestion
digestion
digestion
digestion
digestion
digestion
digestion
1
1
2
3
3
3
3
4
d26Mg*
(&)
±2 se
(&)
d26Mg*
(&)
±2 se
(&)
d25Mg
(&)
±2 se
(&)
d26Mg (&)
±2 se
(&)
n
Bracketing
standard
ChemistryA
DSM-3
J11 olivineB
J11 olivine
DSM-3
J11 olivine
J11 olivine
DSM-3
DSM-3
a
a
a,
a,
a,
a,
a,
a,
J11 olivine
DSM-3
J11 olivine
DSM-3
a
a
a, dga, c
a, dga, c, cMn
J11 olivine
DSM-3
J11 olivine
DSM-3
a
a
a, dga, c
a, dga, c, cMn
DSM-3
J11 olivine
DSM-3
J11 olivine
DSM-3
J11 olivine
a,
a,
a,
a,
a,
a,
Weighted mean & 2 se
0.0069
0.0061
0.0056
0.0051
0.0028
0.0037
0.0036
0.0043
0.0095
0.0157
0.0190
0.0136
0.0145
0.0194
0.0154
0.0123
0.0089
0.0063
0.0053
0.0046
0.0040
0.0067
0.0070
0.0055
0.137
0.054
0.018
0.095
0.060
0.049
0.164
0.156
0.081
0.125
0.033
0.030
0.034
0.050
0.048
0.087
0.281
0.092
0.055
0.198
0.135
0.115
0.340
0.320
0.161
0.250
0.061
0.058
0.067
0.100
0.096
0.172
6
7
9
14
13
7
7
8
-0.0124
0.0172
0.0025
0.0025
0.0131
0.0166
0.0030
0.0026
0.138
0.018
0.031
0.051
0.285
0.053
0.063
0.103
35/4/3
36/4/3
0.0137
0.0080
0.0145
0.0104
0.0049
0.0046
0.0038
0.0040
0.0140
0.0084
0.0145
0.0104
0.0064
0.0069
0.0047
0.0055
0.043
0.201
0.035
0.091
0.111
0.056
0.027
0.066
0.072
0.401
0.054
0.186
0.230
0.113
0.054
0.133
6
9
10
10
-0.0092
0.0141
0.0024
0.0008
0.0096
0.0143
0.0043
0.0038
0.146
0.039
0.110
0.008
0.294
0.063
0.215
0.018
19/2/2
16/2/2
0.0138
0.0112
0.0176
0.0137
0.0061
0.0120
0.0048
0.0038
0.0143
0.0110
0.0173
0.0136
0.0060
0.0072
0.0056
0.0062
0.089
0.121
0.094
0.021
0.088
0.076
0.029
0.081
0.157
0.247
0.202
0.032
0.175
0.143
0.046
0.158
8
9
9
8
-0.0125
0.0157
0.0025
0.0038
0.0125
0.0159
0.0047
0.0041
0.050
0.003
0.142
0.183
0.108
0.023
0.279
0.359
17/2/2
17/2/2
0.0117
0.0136
0.0130
0.0206
0.0105
0.0137
-0.0117
0.0160
0.0035
0.0031
0.0042
0.0034
0.0028
0.0049
0.0014
0.0046
0.0114
0.0131
0.0127
0.0208
0.0106
0.0139
0.0117
0.0169
0.0061
0.0058
0.0056
0.0046
0.0063
0.0064
0.0034
0.0031
0.133
0.065
0.207
0.030
0.050
0.038
0.130
0.044
0.032
0.058
0.019
0.018
0.043
0.038
0.091
0.021
0.272
0.146
0.417
0.073
0.108
0.087
0.266
0.102
0.065
0.113
0.037
0.033
0.086
0.075
0.179
0.045
7
8
11
16
8
10
26/3/3
34/3/3
dga,
dga,
dga,
dga,
dga,
dga,
c
c
c
c, cMn
c, cMn
c, cMn
Admire mean versus DSM-3
Admire mean versus J11
Brenham
Brenham
Brenham
Brenham
digestion
digestion
digestion
digestion
1
1
2
3
Fo87.6
“
“
“
0.000
“
“
“
Brenham mean versus DSM-3
Brenham mean versus J11
Esquel
Esquel
Esquel
Esquel
digestion
digestion
digestion
digestion
1
1
2
3
Fo88.3
“
“
“
0.000
“
“
“
Esquel mean versus DSM-3
Esquel mean versus J11
Molong
Molong
Molong
Molong
Molong
Molong
Molong
Molong
digestion 1
digestion 1
digestion 2
digestion 2
digestion 3
digestion 3
mean versus DSM-3
mean versus J11
Fo88.7
“
“
0.000
“
“
“
“
“
“
dga,
dga,
dga,
dga,
dga,
dga,
Al–26Mg deficit dating ultramafic meteorites
26
Average & 2 se
0.0093
0.0163
0.0191
0.0127
0.0140
0.0194
0.0153
0.0124
c
c
c
c
c, cMn
c, cMn
A
a = Anion chemical separation (Fe); dga = TODGA Eichrom separation (Ca); c = cation exchange separation (most elements except Mn and Ni); cMn = cation exchange separation in HCl/
acetone (Mn).
B
Mg separated from J11 – a mantle olivine from an anhydrous spinel peridotite (Jordan).
421
422
Table 4
Mg isotope and Al/Mg data for ureilite and aubrite meteorites.
Type
27
Al/24Mg
Ureilites
SAH98505
El Gouanem
NWA2234
NWA766
Aubrites
Norton County
Pena Blanca Spring
Mt. Egerton
Shallowater
Cumberland Falls
Bishopville
Bishopville (pyroxene)
A
d26Mg*
(&)
±2 se
(&)
d26Mg*
(&)
±2 se
(&)
Average & 2 se
Weighted mean & 2 se
d25Mg
(&)
±2 se
(&)
d26Mg
(&)
±2 se
(&)
n
Bracketing
standard
ChemistryA
Fo79-83
“
Fo76-78
“
Fo82-92
“
Fo76
“
0.010
“
0.005
“
0.025
“
0.070
“
0.0038
0.0099
0.0048
0.0061
0.0126
0.0091
0.0043
0.0058
0.0120
0.0058
0.0091
0.0077
0.0020
0.0039
0.0070
0.0058
0.0040
0.0090
0.0046
0.0067
0.0124
0.0084
0.0042
0.0060
0.0046
0.0034
0.0045
0.0064
0.0047
0.0046
0.0046
0.0053
0.218
0.024
0.341
0.066
0.2725
0.2075
0.152
0.123
0.277
0.124
0.128
0.166
0.227
0.177
0.128
0.074
0.431
0.037
0.673
0.130
0.5435
0.4015
0.304
0.235
0.551
0.248
0.257
0.331
0.439
0.351
0.260
0.142
22/2/1
38/3/1
20/2/1
9/1/1
20/2/1
19/2/1
22/2/1
16/1/1
DSM-3
J11 olivineB
DSM-3
J11 olivine
DSM-3
J11 olivine
DSM-3
J11 olivine
a,
“
a,
“
a,
“
a,
“
–
–
–
–
–
–
–
0.0017
0.0018
0.0022
0.0046
0.0054
0.370
0.0240
0.0006
0.0030
0.0054
0.0011
0.0022
0.0036
0.0003
0.0052
0.0044
0.0077
0.0060
0.0051
0.0056
0.0087
0.0009
0.0025
0.0056
0.0006
0.0019
0.0036
0.0001
0.0039
0.0054
0.0063
0.0060
0.0045
0.0058
0.0064
0.179
0.119
0.091
0.113
0.060
0.130
0.067
0.049
0.038
0.019
0.018
0.063
0.110
0.023
0.352
0.230
0.175
0.226
0.110
0.260
0.133
0.096
0.074
0.038
0.030
0.130
0.210
0.041
24/2/1
10/1/1
10/1/1
10/1/1
19/2/1
10/1/1
10/1/1
DSM-3
DSM-3
DSM-3
DSM-3
DSM-3
DSM-3
DSM-3
a,
a,
a,
a,
a,
a,
a,
dga, c, cMn, dmg
dga, c, cMn, dmg
dga, c, cMn, dmg
dga, c, cMn, dmg
dga,
dga,
dga,
dga,
dga,
dga,
dga,
c,
c,
c,
c,
c,
c,
c,
cMn
cMn
cMn
cMn
cMn
cMn
cMn
a = Anion chemical separation (Fe); dga = TODGA Eichrom separation (Ca); c = cation exchange separation (most elements except Mn and Ni); cMn = cation exchange separation in HCl/
acetone (Mn); dmg = dimethylgloxamine Ni-specific chemistry.
B
Mg separated from J11 – a mantle olivine from an anhydrous spinel peridotite (Jordan).
J.A. Baker et al. / Geochimica et Cosmochimica Acta 77 (2012) 415–431
Sample
26
Al–26Mg deficit dating ultramafic meteorites
423
26
Mg* excesses of 0.010&, 0.020&, 0.030& and 0.200&.
These solutions were then analysed a number of times versus undoped Aristar Mg. While the doped standards were
analysed a variable number of times with differences in
the resultant 2 se on the final d26Mg* values, all solutions
yielded d26Mg* values within 2 se analytical uncertainties
of the expected value (Table 1 and Fig. 2). For example,
the solution with a 0.010& 26Mg excess produced a mean
d26Mg* = 0.0107 ± 0.0060& after eight analyses.
3.1.3. Aristar Mg subjected to chemical separation of Mg
Approximately 1 mg of Aristar Mg was processed
through two different types of chemistries – anion exchange
separation of Fe and (on two occasions) anion exchange
separation of Fe + TODGA separation of Ca + cation exchange separation of most major and trace elements. Both
the abundance of 26Mg (d26Mg*) and stable Mg isotopic
composition of the column-processed standards produced
values within 2 se analytical uncertainties of zero (Table 1
and Fig. 2).
3.1.4. Aristar Mg collected sequentially from a cation
exchange separation column
Aristar Mg collected sequentially from a cation exchange column in both 1 M HNO3 and 1 M HCl shows that
heavier isotopes of Mg preferentially pass faster through
the column as compared to lighter isotopes of Mg (Table
2). These different Mg cuts do not generally yield calculated
d26Mg* values that are within 2 se analytical uncertainties
of zero. In particular, heavy (d25Mg = +1.38&) and light
(d25Mg = 0.76&) Mg in these cuts is characterised by
apparent deficits (d26Mg* = 0.0315&) and excesses
(d26Mg* = 0.0210&) in the abundance of 26Mg relative to
unprocessed Aristar Mg (Table 2 and Figs. 2 and 3).
3.1.5. Analysis of terrestrial olivine standards
Mg separated from olivine crystals taken from three terrestrial materials was analysed versus DSM-3 and Mg separated from J11 mantle olivine. The mean d26Mg* obtained
on Mg from J11 mantle olivine analysed versus DSM-3 was
+0.0029 ± 0.0028& (Table 1 and Fig. 2). The stable Mg
isotopic composition of d25 MgDSM-3 ¼ 0:10 0:06& is
within error of the value published by Handler et al.
(2009). Mg separated from mantle olivine samples (J11
and JK3) as well as basaltic olivine (JB281) yields d26Mg*
(+0.0005 ± 0.0034&) and d25Mg (0.029 ± 0.062&) values,
as would be expected, within 2 se uncertainty of zero when
measured versus Mg separated from J11 olivine in a previous digestion and chemical separation pass.
3.1.6. Analysis of the Mg isotopic composition of J11 mantle
and pallasite olivine samples after different stages of chemical
separation of Mg
Mg separated from both the J11 mantle and pallasite olivines was measured after different stages (anion exchange
separation of Fe ± TODGA separation of Ca ± cation exchange separation of most major and trace elements ± cation exchange separation of Mn in 0.5 M HCl–95% acetone)
of chemical separation of Mg to assess the extent to which
the progressive removal of matrix elements affected the re-
Fig. 2. Summary of the abundance of 26Mg (d26Mg*) of pallasite
olivine, ureilite and aubrite meteorites, terrestrial standards, and
chondrite meteorites (Schiller et al., 2010a). The grey field
represents the range of terrestrial mantle and basalt olivine.
sults. The results show that consistent results for the d26Mg*
of both J11 and pallasite olivine are produced irrespective
of the extent of the chemical separation procedures used
to purify Mg, provided that Si (HF digestion) and Fe
(anion exchange separation) have been removed from the
olivine (Tables 1 and 3 and Fig. 4). When all data are combined, the average d26Mg* of both the terrestrial and pallasite olivines are marginally more positive (0.0038 ±
0.0021&) when DSM-3 is used as the bracketing standard
rather than Mg separated from J11 mantle olivine (Tables
1 and 3).
424
J.A. Baker et al. / Geochimica et Cosmochimica Acta 77 (2012) 415–431
26
26
*
Fig. 3. Variations in the abundance of Mg (d Mg ) in isotopically fractionated (d25Mg) cuts of Mg produced by sequentially
collecting Mg from cation exchange columns (AG50W-X8 resin).
The line on the graph represents the expected (erroneous) d26Mg*
values that would result from applying a kinetic/exponential mass
bias (b = 0.511) correction to the isotopic analyses when, in fact,
Mg has been fractionated on the cation exchange columns by an
equilibrium (b = 0.521) fractionation process.
Fig. 4. d26Mg* values for terrestrial olivine (J11, JK3, JB281) and
pallasite olivine (Admire) analysed after different chemical separation procedures and utilising different bracketing standards (i.e.,
DSM-3 [filled square symbols] and Mg separated from J11 mantle
olivine [open square symbols]). a = anion exchange separation of
Fe; c = anion exchange separation of Fe + TODGA separation of
Ca + cation exchange separation of most major and trace elements
(except Mn and Ni); Mn = anion exchange separation of Fe +
TODGA separation of Ca + cation exchange separation of most
major and trace elements + cation exchange separation of Mn in
0.5 M HCl–95% acetone.
3.2. Pallasites
All of the olivine separates from the main group pallasites
have 27Al/24Mg ratios that are effectively zero (<0.001).
Analyses of the abundance of 26Mg of the pallasite olivines
repeatedly show resolvable deficits with respect to the terrestrial standard, whether bracketed by analyses of DSM-3 or
Mg separated from J11 mantle olivine (Table 3 and
Fig. 2), and irrespective of the chemical separation methodology used to purify Mg (Fig. 4). The mean d26 MgDSM-3 of all
the pallasite olivine analyses is 0.0120 ± 0.0018&, which is
Fig. 5. 26Al–26Mg isochron diagram for pallasite olivine and
ureilite meteorites. Isochrons are anchored with the mean values
determined for non-CAI bearing chondrites (Schiller et al., 2010a;
27
Al/24Mg = 0.091 and d26Mg* = 0.0015 ± 0.0013&; 2 se). Initial
26
Al values are calculated using our conservative estimate of data
reproducibility i.e., 1.5 times the 2 se uncertainty and only using
data obtained versus the DSM-3 standard. Also shown is the
regression through data for chondrite meteorites and CAIs
(Schiller et al., 2010a).
slightly less negative than the value obtained when pallasite
olivines were measured against Mg separated from J11 mantle olivine i.e., d26 MgJ11 ¼ 0:0162 0:0016&. There is no
resolvable difference in d26Mg* values between olivines from
the different main group pallasites. We use a model approach
to calculate initial 26Al/27Al ratios of pallasites (and other
meteorites) studied here by assuming that each meteorite
group originated from a parent body that accreted from
material with Al/Mg ratios and a present-day Mg isotope
composition represented by non-CAI bearing chondrite
meteorites. A regression through the 27Al/24Mg – d26Mg*
pallasite olivine data including the average composition of
non-CAI-bearing chondrites (Schiller et al., 2010a) defines
a line with a slope and initial (26Al/27Al) = 1.6 ± 0.5 105
(Fig. 5).
The stable Mg isotopic composition of olivine from the
four main group pallasites show very limited variations
from d25MgDSM-3 = 0.05 ± 0.14& to 0.15 ± 0.11&,
which are within error of the data previously reported by
Handler et al. (2009) for pallasite olivine, and these overlap
the values for olivine from Earth’s upper mantle. These values are also identical to those reported by Teng et al. (2010)
for a wide range of oceanic basalts, peridotite xenoliths and
chondrite meteorites. Our d25MgDSM-3 values for pallasite
and mantle olivine are, however, inconsistent with those
published by Chakrabarti and Jacobsen (2010) whose
d25MgDSM-3 values for pallasite and mantle olivine are systematically lighter due to as yet unknown analytical artefacts that have likely comprised the accuracy of Mg stable
isotopic data presented by Chakrabarti and Jacobsen
(2010).
3.3. Ureilites
Three of the ureilites yielded markedly sub-chondritic
Al/24Mg ratios from 0.005 to 0.025 (El Gouanem,
SAH98505, NWA2234) (Table 4). However, NWA766
27
26
Al–26Mg deficit dating ultramafic meteorites
has a higher 27Al/24Mg (0.070) that may be consistent with
the presence of high Si–Al glass in this ureilite.
All d26Mg* values measured in the ureilites whether
bracketed by DSM-3 or Mg separated from J11 olivine
yield slight deficits in the abundance of 26Mg relative to
the terrestrial standards, although in some cases
(SAH98505 and NWA766 bracketed by DSM-3) these are
just within 2 se analytical uncertainty of the terrestrial standard (Table 4 and Fig. 2). The mean d26 MgDSM-3 of all the
ureilite analyses is 0.0062 ± 0.0023& which is slightly less
negative than when these samples were measured against
Mg separated from J11 mantle olivine i.e., d26 MgJ11 ¼
0:0080 0:0023&. The d26 MgDSM-3 measured on ureilites
in this study are within analytical uncertainty of those measured by Larsen et al. (2011) on two ureilites, including
SAH98505. A regression through the 27Al/24Mg – d26Mg*
ureilite data including the average composition of nonCAI-bearing chondrites (Schiller et al., 2010a) defines a line
with a slope and initial (26Al/27Al) = 8.8 ± 7.7 106
(Fig. 5). Stable Mg isotope data range from d25MgDSM3 = 0.15 ± 0.13& to 0.34 ± 0.13& and overlap values
measured for samples of Earth’s mantle and basalts as well
chondrites (Handler et al., 2009; Yang et al., 2009; Schiller
et al., 2010a; Teng et al., 2010).
3.4. Aubrites
With the exception of Bishopville, the six bulk aubrite
samples all have markedly sub-chondritic 27Al/24Mg ratios
that range from 0.0017 (Norton County) to 0.0054 (Cumberland Falls) (Table 4). Bishopville has a 27Al/24Mg ratio
(0.370) that is considerably higher than that measured on
a larger bulk sample of this meteorite (0.039; Easton,
1985) suggesting that a feldspar-rich region of Bishopville
was initially sampled in this study. Therefore, a second
pyroxene-rich mineral separate of this meteorite was subsequently prepared, digested and analysed and yielded a lower 27Al/24Mg = 0.0240.
The abundance of 26Mg for all the aubrite samples,
including the high 27Al/24Mg sample of Bishopville, are
all within 2 se analytical uncertainties of the terrestrial standard used to bracket the analyses (DSM-3) (Table 4 and
Fig. 2). The mean d26 MgDSM-3 of all the aubrite analyses
is +0.0015 ± 0.0020&. A regression through the 27Al/24Mg
– d26Mg* data including the average composition of nonCAI-bearing chondrites (Schiller et al., 2010a) defines a line
with a slope of zero and maximum possible slope and initial
(26Al/27Al) = 3.3 106 (given the error on the regression).
Stable Mg isotope data range from d25MgDSM-3 =
+0.13 ± 0.11& to 0.18 ± 0.05& and overlap values measured for samples of Earth’s mantle and basalts (Handler
et al., 2009; Yang et al., 2009; Teng et al., 2010; Schiller
et al., 2010a).
4. DISCUSSION
4.1. Precision and accuracy of d26Mg* data
Resolving small deficits in 26Mg abundances in meteorites for dating purposes requires a careful assessment as
425
to whether it is possible to accurately and precisely measure
d26Mg* values to <±0.005&. Multiple Mg isotope analyses
(n = 10–40) of single samples, by pooling of analyses, which
is a common practice in application of all short-lived chronometers to early Solar System chronometry (e.g., Lugmair
and Shukolyukov, 1998; Kleine et al., 2005a,b; Markowski
et al., 2006a,b; Wadhwa et al., 2009; Villeneuve et al., 2010;
Bizzarro et al., 2011), can produce weighted mean d26Mg*
values with internal 2 standard errors (se) as low as
±0.006& to 0.002& (Tables 1–4).
The 2 se values quoted in this study include the uncertainties incorporated from the bracketing standards as well
as that on the sample. Previous high precision Mg isotope
studies of meteoritic material (e.g., Baker et al., 2005; Bizzarro et al., 2005) did not calculate d26Mg* as weighted
means or incorporate uncertainties from the bracketing
standard (or the sample) into the final 2 se and simply calculated the average and 2 se from the d26Mg* values as the
mean of n measurements and 2 se = 2 sd/n. The difference
in doing this (“average & 2 se”) as compared to the approach adopted here (“weighted mean & 2 se”) and in Schiller et al. (2010a,b) is evident from Tables 1, 3 and 4. While
the two different approaches do not result in significant differences in the mean or average d26Mg* values of more than
0.0015&, it is common for the quoted 2 se to vary significantly using the two different approaches. In about 55%
of cases the 2 se calculated without incorporating the uncertainties from the sample and standard analyses (“average
and 2 se”) is lower than that calculated using the weighted
mean approach. In 30% of cases the calculated 2 se values
are essentially the same (within 0.001&) irrespective of
the method used to calculate them. We consider that the approach adopted herein provides a more realistic and conservative estimate of the analytical uncertainties on Mg
isotope analysis by MC-ICPMS as the “average and 2 se”
approach can yield overly optimistic estimates of 2 se, particularly where the number of repeat analyses (n) is small
and, fortuitously, a small spread in individual d26Mg* values is obtained.
The accuracy of the presented Mg isotope data can be
assessed by the analyses of column-processed standards,
gravimetrically 26Mg spiked standards, terrestrial olivines
and terrestrial and pallasite olivines subject to different
amounts of chemical purification of Mg. In all cases, these
data yield d26Mg* values that are, at worst (J11 digestion 3;
Table 1), within 1.4 times the 2 se of the expected value and,
apart from this example, all within 2 se uncertainties of the
expected values. This demonstrates that the data are accurate to quoted uncertainties i.e., <1.5 times the final 2 se obtained on any particular sample.
While it is possible to measure d26Mg* both precisely
and accurately to < ± 0.005&, our analytical tests do reveal
some artefacts that have potential to produce inaccurate
data, although these are not important for the meteorite
data obtained in this study. In particular, some isotopically
fractionated and light ICP-MS Mg solution standards (Alfa
Aesar and Aristar) have apparent excesses in d26Mg*. This
reflects an artefact of these standards containing Mg that
has, in part, experienced equilibrium stable isotopic fractionation. When an exponential/kinetic mass fractionation
426
J.A. Baker et al. / Geochimica et Cosmochimica Acta 77 (2012) 415–431
law (b = 0.511) is applied to these isotopically light standards to correct the d26Mg* to the d25Mg values of DSM3 it results in erroneously positive d26Mg*. This effect has
been previously noted, described and accounted for by Bizzarro et al. (2011) and, in particular, it is noteworthy that
two Alfa Aesar Mg solutions these authors analysed versus
the DSM-3 standard were both characterised by isotopically light Mg (i.e., d25Mg) and positive d26Mg* values of
ca. 0.02&, which are within error of the d26Mg* values obtained for the two Alfa Aesar Mg solutions analysed in this
study (Table 1).
A similar effect is evident from the Mg that was deliberately isotopically fractionated on large resin bed cation exchange columns (Fig. 3). Here isotopically light and heavy
Mg have apparent excesses and deficits, respectively, in
d26Mg* as the fractionation processes on the cation exchange resin appear to be dominated by an equilibrium process. d26Mg* values re-calculated for all but two of the
heaviest Mg cuts using the equilibrium mass fractionation
law (b = 0.521) yield d26Mg* values within error of
0.000&. Interestingly, the two heaviest cuts, which pass
through the columns first appear to also have been affected
by a component of kinetic and equilibrium isotopic fractionation, suggesting that Mg close to the chromatographic
“front” is not undergoing a purely equilibrium isotopic fractionation with the cation exchange resin. These results highlight the need to carefully chose an appropriate standard
and ensure nearly 100% Mg yields for high precision Mg isotope studies. Similar stable isotope fractionations of other
elements like Li and Ca on ion exchange columns have been
observed (Lee and Begun, 1958; Russell and Papanastassiou, 1978), but the data presented here also shows for the
first time that mass-independent isotopic abundances can
be affected by these processes as well, and must be considered in all ultra-precise studies of isotopic anomalies.
Despite these observations, the meteorite data presented
in Tables 3 and 4 are unaffected by these analytical artefacts
as a bracketing standard with similar stable Mg isotope
composition (DSM-3 or Mg separated from J11 olivine)
was used in obtaining these results, and because the chemical yield of Mg was >99% irrespective of whether a onestep anion chemical separation of Mg was used or the full
five-step chemical separation.
One further aspect of the d26Mg* data presented in Tables 1 and 3 where data for samples was obtained by measurement against DSM-3 and also Mg separated from J11
mantle olivine is that d26Mg* values measured versus
DSM-3 are consistently, but only marginally, more positive
(0.0038 ± 0.0021&) than when measured versus Mg separated from J11 mantle olivine. This might reflect the fact
that DSM-3 has experienced a small amount of equilibrium
stable isotopic fractionation (0.1&) as compared to the
bulk Earth stable Mg isotopic composition as represented
by analyses of upper mantle olivine (Handler et al., 2009).
Alternatively, the presence of very small amounts of Ni in
the Mg cuts from the terrestrial olivine may be responsible
for these small differences. These differences are near the
resolution of our analytical uncertainties, but we consider
that the d26Mg* values measured versus DSM-3 are the
most accurate data in this study, given that an ultra-precise
Mg isotope study of terrestrial solution, rock and mineral
standards by Bizzarro et al. (2011) did not observe small
positive d26Mg* values for these standards when measured
against DSM-3.
4.2. Do the d26Mg* deficits in pallasite olivines and ureilites
have chronological significance?
Before interpreting the measured 26Mg* deficits in pallasite olivines and ureilites as having chronological significance it is necessary to consider whether these could be
caused by non-analytical artefacts such as nuclear reactions
due to exposure to cosmic rays or heterogeneous distribution of 26Al and/or Mg isotopes on a planetesimal scale
in the Solar System.
Mg isotopes have small thermal neutron capture cross
sections (0.054–0.200 barn; Walkiewicz et al., 1992) and
coupled with the relatively short cosmic ray exposure ages
of the studied classes of meteorites (ca. 0–200 Myr; Eugster,
2003; Herzog, 2003) means that it is unlikely that cosmogenic effects are responsible for the small variations in
d26Mg*. It is also notable that the aubrite Norton County,
which has amongst the longest cosmic ray exposure age
(112 Myr) of stony meteorites, has a d26Mg* value identical
to the other aubrites, which have shorter cosmic ray exposure ages.
26
Al or Mg isotope heterogeneity on a planetesimal scale
in the Solar System as has been documented for neutronrich isotopes of Ti, Cr and Ni (Trinquier et al., 2007,
2009; Regelous et al., 2008) might also be potentially invoked as being responsible for the observed variations in
d26Mg*. However, high precision Mg isotope analysis of
bulk chondrites has failed to detect significant Mg isotopic
heterogeneity beyond that due to the presence of CAIs in
some carbonaceous chondrites, and does not hint at
large-scale 26Al or Mg isotope heterogeneity on a planetesimal scale in the proto-planetary disc (Schiller et al., 2010a),
although this study concluded it was not possible to establish if planetesimals accreted from material that initially
had the same levels of 26Al as CAIs. A range of nonCAI-bearing chondrites have a mean d26Mg* = 0.0015 ±
0.0013&. Thus, the d26Mg* deficits observed in pallasites
and, potentially, ureilites are most likely the result of the
lack of in-growth of 26Mg due to development of low Al/
Mg ratios as a result of silicate planetary differentiation
very early in the Solar System.
4.3. Silicate differentiation ages for pallasites, ureilites and
aubrites
In the first instance, model initial 26Al/27Al values can be
calculated for the pallasite olivines, ureilites and aubrites by
regressing the 27Al/24Mg – d26Mg* data for each class of
meteorites with the mean 27Al/24Mg – d26Mg* values for
non-CAI-bearing chondrites (Fig. 5). The pallasite olivines
and ureilites both yield non-zero abundances of initial
26
Al/27Al of 1.6 ± 0.5 105 and 8.8 ± 7.7 106. However, the aubrite data yields a regression with a slope of zero
and a potential maximum initial 26Al/27Al of 3.3 106, given the uncertainty on the regression.
26
Al–26Mg deficit dating ultramafic meteorites
We first convert these to relative ages with respect to
CAIs using an initial 26Al/27Al of 5.21 105 based on a
regression of high precision Mg isotope data for CAIs (Jacobsen et al., 2008) with bulk chondrites (Schiller et al.,
2010a). Given recent documentation of 238U/235U variations in CAIs (Brennecka et al., 2010) we do not convert
these to absolute ages using Pb–Pb ages for CAIs. Subsequently, we consider the validity of these model ages, if
26
Al was heterogeneously distributed in the young Solar
System as recently suggested by Larsen et al. (2011).
4.3.1. Pallasite olivine ages
The pallasite data result in a relative age for pallasite
olivine of 1:24þ0:40
0:28 Myr after CAI formation. This age is a
model date for formation of sub-chondritic and near-zero
Al/Mg in pallasite olivine and records the last time the olivine was in equilibrium with parts of their parent body with
chondritic or super-chondritic Al/Mg. In the case of the
pallasite olivine, this presumably reflects crystallization of
olivine and diffusive isolation at the base of silicate magma
oceans near the core-mantle boundary of a planestesimal
formed early in Solar System. Simple thermal models for
accretion of a differentiated planetesimal (e.g., Bizzarro
et al., 2005) heated by 26Al decay would imply from the
pallasite olivine model age that the main group pallasite
parent body accreted within 0.3 Myr of CAI formation.
The age constraints presented here from high precision
Mg isotope data for main group pallasites are in agreement
with other available chronological data for pallasites,
although this group of meteorites has not been clearly
and precisely dated. Precise age constraints on pallasites
are scarce but Hf–W isotope data for some pallasite metals
(Quitté et al., 2005) overlaps the initial e182W value of CAIs
suggesting very early formation that coincided with formation of magmatic iron meteorites and CAIs. Recently, Villeneuve et al. (2010) presented in situ Mg isotope data for the
Eagle Station pallasite olivine, which had a large measured
d26Mg* deficit of -0.033 ± 0.008&, which corresponds to an
earlier model age of 0:15þ0:29
- 0:23 Myr after CAI formation
than the age for main group pallasite olivines presented
here. While this is generally consistent with the very early
silicate differentiation age for pallasite olivine presented
here, in detail the different ages might represent one of
two factors: (1) The Eagle Station parent body which has
a distinct mass-independent oxygen isotope composition
as compared to the main group pallasite parent body accreted and differentiated earlier than the main group pallasite parent body; (2) the in situ Mg isotope data for Eagle
Station olivine are potentially inaccurate due to analytical
artefacts.
4.3.2. Ureilite bulk rock ages
The ureilite data results in a relative age for ureilites of
þ2:2
1:90:7
Myr after CAI formation. The ureilite model age
most likely dates silicate partial melting and/or smelting
of primitive material and extraction of a basaltic component with super-chondritic Al/Mg from the silicate “mantle” of a planetesimal. The age constraints presented here
from high precision Mg isotope data of ureilites are in
agreement with other available chronological data for
427
ureilites. However, again, this group of meteorites has yet
to be clearly and precisely dated although Torigoye-Kita
et al. (1995b) reported a U–Pb age of 4.563 ± 0.006 Ma
for the ureilite MET 78008. Bulk ureilites have e182W values
and sub-chondritic Hf/W ratios that generally overlap the
initial e182W value of CAIs, which has been interpreted as
reflecting differentiation of the ureilite parent body within
1–2 Myr after the start of the Solar System (Lee et al.,
2009). Goodrich et al. (2010) also presented in situ
53
Mn–53Cr and 26Al–26Mg mineral isochron ages for feldspathic clasts in two polymict ureilites, which yielded relative ages of 4–5 Myr after CAI formation. These ages
overlap and/or are slightly younger than the Mg isotope
model ages for silicate melting of ureilites presented here.
The two types of ages are in excellent agreement given that
the Mg isotope model ages date silicate melting of the ureilite parent body, which must pre-date the crystallization of
the feldspathic clasts in the polymict ureilites as measured
by in situ mineral 53Mn–53Cr and 26Al–26Mg isochrons
from the type of melt with super-chondritic Al/Mg extracted from monomict ureilites as dated by the high precision bulk Mg isotope data. Relatively late accretion and
silicate differentiation of the ureilite parent body as compared to the pallasite parent body is consistent with the evidence that this planetesimal did not undergo complete
melting, principally due to the lower levels of 26Al available
to generate melting through radioactive decay of this shortlived isotope due to later accretion of the ureilite parent
body.
4.3.3. Aubrite ages
Aubrites can only be constrained to have formed at least
2.9 Myr after CAI formation. In the case of aubrites this
presumably reflects crystallization of pyroxene and diffusive
isolation at the base of a silicate magma ocean or in a magma chamber within the aubrite parent body. Aubrite meteorites have yielded few reliable and precise age constraints,
probably due to their complex thermal history, but the age
constraints for their formation presented here (>2.9 Myr
after CAIs) is consistent with Mn–Cr isotope data for aubrites (Shukolyukov and Lugmair, 2004). Petitat et al. (2008)
measured 182Hf–182W metal–silicate isochrons for aubrites,
which yielded two groups of ages of 4550 and 4560 Ma,
which are generally consistent with the age data presented
here. However, the Mg isotope model ages and Hf–W model ages likely date different events, namely formation of the
aubrite pyroxene cumulates (Mg) and internal requilibration of metal and silicate (Hf–W) in the aubrite samples
and, as such, are not directly comparable.
It should be noted that these 26Al–26Mg deficit model
ages discussed above are only accurate if planetesimals in
the proto-planetary disc accreted from material that
initially had the same levels of 26Al as CAIs. Schiller
et al. (2010a) concluded from a high precision Mg isotopic
study of chondrites that it is not possible with the currently
available data to determine with certainty whether CAIs
and the material from which planetesimals accreted had
precisely the same initial levels of 26Al, with the proto-planetary disc potentially having ca. 40–130% of the levels of
26
Al in CAIs. Thus, these relative ages for silicate
428
J.A. Baker et al. / Geochimica et Cosmochimica Acta 77 (2012) 415–431
differentiation after CAI formation on planetesimals may
be too long (by ca. 1.0 Myr) or too short (by ca. 0.4 Myr)
given these constraints on whether the initial 26Al abundance in the proto-planetary disc was present on a planetesimal scale at levels equivalent to CAIs based on this
comparison of high precision Mg isotope data for CAIs
and bulk chondrites (Jacobsen et al., 2008; Schiller et al.,
2010a). Moreover, a recent high-precision 26Al–26Mg isochron for CAIs and amoeboid olivine aggregates (Larsen
et al., 2011) coupled with high precision Mg isotope data
for a range of planetary materials and 54Cr isotopic data
has been used to argue that significant 26Al heterogeneity
may have existed in the early Solar System with the planet
and planetesimal-forming region having markedly (up to
ca. 80%) lower levels of 26Al than the CAI-forming region
of the proto-planetary disc. While the degree of 26Al heterogeneity inferred by Larsen et al. (2011) differs slightly from
that suggested to be possible by Schiller et al. (2010a), if the
conclusions of Larsen et al. (2011) are correct, then the silicate differentiation ages for ureilites and pallasite olivines
may need to be reconsidered. In the case of ureilites, the
lower initial 26Al in their parent body (1.1 ± 0.3 105)
suggested by Larsen et al. (2011) means that no age constraints can be placed on the timing of their silicate melting
and differentiation. The initial 26Al and d26Mg* of the main
group pallasite parent body can be estimated from the 54Cr
isotopic composition of pallasites (Trinquier et al., 2007),
which is intermediate between that of angrites and ureilites,
and the correlation between d26Mg* and 54Cr shown by
Larsen et al. (2011), as being 1.3 105 and 0.006&,
respectively. Recalculating the model initial 26Al/27Al
values for pallasite olivines using chondritic Al/Mg
and d26 Mginitial ¼ 0:006& for its parent body yields
0.9 ± 0.5 105, which corresponds to a relative age of
Fig. 6. 26Mg deficit ages for silicate differentiation of planetesimals
in the early Solar System compared with selected 26Al–26Mg and
182
Hf–182W ages relative to CAIs for: (a) metal-silicate fractionation and core formation on planetesimals as represented by
magmatic iron meteorites with low cosmic ray exposure ages
(Burkhardt et al., 2008), and asteroidal basaltic magmatism
(Bizzarro et al., 2005; Markowski et al., 2007; Burkhardt et al.
2008; Spivak-Birndorf et al., 2009; Schiller et al., 2010b) on
achondrite or differentiated planetesimals (angrites, eucrites and
NWA2976); (b) chondrule formation (Amelin et al., 2002; Krot
et al., 2005; Villeneuve et al., 2009).
0:4þ0:9
0:5 Myr after CAI formation. Estimation of the initial
Al/27Al values for pallasite olivines using this approach
adds an additional source of uncertainty to the model
age, which is difficult to quantify. However, irrespective
of current estimates of the initial 26Al and d26Mg* of the
main group pallasite parent body it is apparent that
d26Mg* deficits date its silicate differentiation to the first
1.5 Myr of the Solar System.
The 26Al–26Mg deficit ages for silicate differentiation on
the pallasite and ureilite parent bodies are shown in Fig. 6.
Irrespective of the assumed initial 26Al abundance present
on a planetesimal scale in the proto-planetary disc, the ages
for silicate differentiation on the main group pallasite parent body are intermediate between those for metal-silicate
fractionation for core formation obtained from magmatic
iron meteorites (e.g., Kleine et al., 2005b; Markowski
et al., 2006a,b; Schersten et al., 2006; Burkhardt et al.,
2008) and those for asteroidal silicate magmatism (e.g., Bizzarro et al., 2005; Markowski et al., 2007; Spivak-Birndorf
et al., 2009; Schiller et al., 2010b; Larsen et al., 2011). This
provides further evidence that differentiated planetesimals
accreted and melted very early in the Solar System, prior
to accretion of undifferentiated planetesimals (chondrites).
26
5. CONCLUSIONS
A high-precision Mg isotope study of ultramafic meteorites (pallasite olivine, ureilites, and aubrites) with sub-chondritic Al/Mg has shown that:
(1) Variations in the abundance of 26Mg (d26Mg*) relative to Earth can be determined in these Mg-rich
materials to 60.005&. However, care needs to be
taken to ensure that yields from chemical separation
of Mg are close to 100% and that an appropriate
standard is used to bracket the Mg isotope analyses
as equilibrium stable isotope effects can generate
small analytical artefacts on the d26Mg* when this
is calculated using the kinetic/exponential mass fractionation law.
(2) Pallasite olivines have clearly resolvable deficits in
d26Mg* (mean d26 MgDSM-3 ¼ 0:0120 0:0018&),
whereas ureilites have smaller and only just resolvable deficits in the abundance of 26Mg (mean
d26 MgDSM-3 ¼ 0:0062 0:0023&), and aubrites
have d26Mg* identical to the terrestrial standards
(mean d26 MgDSM-3 ¼ þ0:0015 0:0020&).
(3) Assuming that 26Al was uniformly distributed
throughout the proto-planetary disc of the young
Solar System, ages can be calculated from these Mg
isotope data that suggest pallasite olivine formed
and was diffusively isolated on the pallasite parent
body 1:24þ0:40
0:28 Myr after CAI formation, whereas
ureilites experienced silicate partial melting
1:9þ2:2
0:7 Myr after CAI formation. Aubrites formed
>2.9 Myr after CAI formation. However, these ages
are highly dependent on the assumption that 26Al
was uniformly distributed. If 26Al was heterogeneously distributed as suggested by Larsen et al.
(2011), then no age constraints can be placed on
26
Al–26Mg deficit dating ultramafic meteorites
silicate melting of ureilites. However, pallasite olivine
d26Mg* deficits still contain a component related to
the absence of in-growth of 26Mg from decay of
26
Al in the early Solar System, yielding a model age
of 0:4þ0:9
0:5 Myr after CAI formation.
(4) Irrespective of the assumed initial 26Al abundance
present on a planetesimal scale in the proto-planetary
disc, the ages for silicate differentiation on the main
group pallasite parent body are intermediate between
those for metal-silicate fractionation for core formation obtained from magmatic iron meteorites (e.g.,
Kleine et al., 2005b; Markowski et al., 2006a,b;
Schersten et al., 2006; Burkhardt et al., 2008) and
those for asteroidal silicate magmatism (e.g., Bizzarro et al., 2005; Markowski et al., 2007; Spivak-Birndorf et al., 2009; Schiller et al., 2010b; Larsen et al.,
2011). Clear evidence now exists that differentiated
planetesimals accreted and melted very early in the
Solar System, prior to accretion of undifferentiated
planetesimals (chondrites).
(5) Identification of d26Mg* deficits in meteoritic material in this study suggests this dating method could
potentially be extended to other types of low Al/
Mg meteorites and their mineral phases such as bulk
samples of diogenites, lodranites and brachinites,
although this will depend on the extent to which
26
Al was homogeneously distributed in the protoplanetary disc.
ACKNOWLEDGEMENTS
Monica Handler is thanked for assistance in VUW’s Geochemistry Laboratory. This work was supported by a VUW URF award
and New Zealand’s Marsden Fund (06-VUW-076). Benjamin
Jacobsen, two anonymous reviewers and the journal AE (Frederic
Moynier) are thanked for their constructive comments on an earlier
version of this paper.
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Associate editor: Frederic Moynier