larkman nunatak 06319 and 12011 enriched olivine-phyric ... · lar 12011, have 60% dark brown to...
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Larkman Nunatak 06319 and 12011 Enriched olivine-phyric shergottites
78.57 and 701.17 grams
Figure 1: LAR 06319 (left) and LAR 12011 (right) found in Larkman Nunatak (LAR) area in the
2006-2007 and 2012-2013 ANSMET seasons, respectively.
Introduction
An olivine-phyric shergottite from Mars, Larkman Nunatak 06319 was reported in the
Antarctic Newsletter 30, #2, 2007 (Figure 1). Larkman Nunatak is in the eastern
Grosvenor Mountains and is the southernmost nunatak in the Shackleton Glacier drainage
(Figure 2 and 3). The site consists of a main nunatak with a few smaller satellite
outcroppings, and the main area of exposed blue ice is north of the largest nunatak with
bare ice areas extending northward to Hayman Nunataks. A reconnaissance team spent
12 days at Larkman Nunatak during the 2004-2005 ANSMET season, and recovered 79
meteorites. ANSMET returned during the 2006-2007 season for systematic search efforts
that resulted in the recovery of 622 meteorite specimens. Another olivine-phyric
shergottite was recovered in the 2012-2013 ANSMET season, LAR 12011, and was
quickly recognized to be paired with LAR 06319 (Righter and Satterwhite, 2013).
According to Righter and Satterwhite (2007; 2013), the exterior of both LAR 06319 and
LAR 12011, have 60% dark brown to black fusion crust with a very fine grained
wrinkled texture (Figures 4 and 5). The fusion crust exhibits a slight sheen. The interior
is a gray and black matrix that is fine grained and very hard (Figures 6 and 7).
Petrography
Thin section descriptions have been published by McCoy et al. (AMN 30, #2, 2007;
AMN 36, #2, 2013; Figure 7), Mittlefehldt and Herrin (2008), Sarbadhikari et al. (2009)
and Shafer et al. (2009). Shafer et al. (2009) describe LAR 06319 as “an olivine-
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Figure 2 (left): an enlarged portion of the U.S.G.S. 1:2500000 scale Plunkett Point map.
Figure 3 (right): oblique aerial view of the Larkman Nunatak area from the south. The Roberts
Massif and Shackleton Glacier are in the distance.
Figure 4 (top): LAR 06319 main mass in the Antarctic Meteorite Lab at NASA JSC.
Figure 5 (bottom): LAR 12011 main mass in the Antarctic Meteorite Lab at NASA JSC.
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Figure 6: Interior images of LAR 06319 (left) and LAR 12011 (right); 1 cm cube for scale.
phyric shergottite consisting of olivine phenocrysts (up to 3 mm) set in a matrix of
pyroxene and maskelynite interspersed with minor oxide phases, phosphate phases and
shock melt veins (Figure 8). The olivine and pyroxene are highly zoned (Figures 9,10).
Maskelynite laths in the matrix are An51-37. Olivine is brownish in color. Chromite, Ti-
chromite, apatite, merrillite, troilite, pyrrhotite and pyrite are also reported (Sarbadhikari
et al., 2009).
Figure 7: Plane polarized light (left) and cross polarized light (right) images of LAR 06319 (top)
and LAR 12011 (bottom), from AMN 30, #2, 2007; AMN 36, #2, 2013
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Sarbadhikari et al. (2009) determined the modal mineralogy and gave detailed mineral
compositions (Table 1). There are both high-Ca and low-Ca pyroxenes similar to
“lherzolitic shergottites” and there are abundant “melt inclusions” in the olivine and
pyroxene cores.
Table 1: Modal mineralogy for LAR 06319 Sarbadhikari 2009
Olivine 24.4 vol. %
Pyroxene
Opx 4.1
Pigeonite 27.7
Augite 22.2
Plagioclase 17.8
Phosphate 2.1
Oxides 1.3
Sulfides 0.3
Figure 8 (left): Photomicrograph of LAR 06319, 37 thin-section in transmitted light; megacrystic
olivine and prismatic pyroxenes are set in a groundmass of smaller olivine (ol), pyroxene (px),
maskelynite (mask), phosphate and spinel grains. Also note the impact melt vein running through
the sample (from Sarbadhikari et al., 2009).
Figure 9 (right): Pyroxene quadrilateral diagram from analyses reported in Mittlefehldt and
Herrin (2008), Sarbadhikari et al. (2009) and Shafer et al. (2009).
Oxides range in composition from chromite to Ti-chromite (Figure 11; Peslier et al.,
2010). The oxides are used together with the pyroxene and olivine to constrain fO2 (see
Petrology section). In addition to the main silicate and oxide phases, there are minor
amounts of the phosphates apatite and merrillite and whitlockite (Howarth et al., 2016;
Peslier et al., 2010; Figure 12). The apatites are dominantly OH-rich (calculated by
stoichiometry) with variable yet high Cl contents (Figure 12). Similarly, Bellucci et al.
(2017) measured halogen (Cl, F, Br, and I) abundances in phosphates from eight Martian
meteorites including LAR (Figure 13) by SIMS and found that they range over several
orders of magnitude - up to some of the largest concentrations yet measured in Martian
samples or on the Martian surface, and the inter-element ratios are highly variable (see
also Figure 14).
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Takenouchi et al. (2018) studied the brown olivine in LAR 06319 and concluded that it
formed at high PT conditions associated with shock metamporphism. They propose that
brown olivine in shergottites is formed by stacking of lamellae as reported in NWA 1950
(Takenouchi et al. 2017). Brown olivine experienced high pressure, high peak
temperature (1750–1870 K), and high postshock temperature (>1200–1170 K) induced
by a shock event. Peak shock pressure may be ~55 GPa, which induces postshock
temperature around ~1270 K (Fritz et al. 2005). Martian meteorites such as LAR 06319,
with brown olivine, also contain no high-pressure phases, suggesting that high postshock
temperatures induced back-transformation of high pressure minerals.
Figure 10: (a) Optical photomicrograph of LAR 06319 ,36 (plane-polarized light). Dark-brown
phenocrysts scattered throughout the section are olivine. Red outline is the x-ray mapped area.
(b) Mg x-ray map. Red-yellow phases are olivine. Pyroxenes are green-blue. Most of the dark
areas correspond to maskelynite. (c) Fe x-ray map. Olivine is shown as green-red areas,
exhibiting extensive chemical zoning. Blue areas are pyroxene. Most dark areas are maskelynite.
(d) Blue is either low-Ca pyroxene or maskelynite. Augite is shown as green, rimming the low-Ca
pyroxene cores. Most dark areas are olivine. (Figure from Park et al., 2013).
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Figure 11: Illustrations of LAR 06319 Cr–Fe–Ti oxide features and compositional variation. (a
and b) Back scattered electron (BSE) image and FeO elemental map (wt%, ZAF corrected)
showing three types of Cr–Fe–Ti oxides. T1 is a homogeneous chromite-rich spinel enclosed in a
pyroxene (Px) megacryst. T1-2-3 is a zoned spinel ranging from Cr-rich and Fe-poor (T1) next to
the pyroxene megacryst core, to Cr-poor and Fe-rich (ulvospinel T2 and T3) towards the
maskelynite (Ma; details shown in c). T5-6 is an assemblage of ilmenite (T5 with 45 wt% FeO)
and magnetite (T6), shown in detail in (d), and sulfides (Sulf). (e) BSE image of an assemblage of
Fe-rich T3 ulvospinel with ilmenite (T4 with 47 wt% FeO) and sulfide. (f) Magnetite
(Fe3+/(Fe3+ + 2Ti + Cr + Al)) vs. ulvospinel (2Ti/(2Ti + Cr + Al)) content for spinels. Also
shown is melt inclusion-hosted spinels (MI). (g) Cr# (Cr/(Cr + Al)) versus Fe# (Fe2+/(Fe2+
+Mg)) of LAR 06319 T1, T2 and T3 spinels. The wide range of Fe# of the T1 spinels reflects in
part the various amounts of Fe–Mg exchange with their host minerals during subsolidus
reequilibration: limited exchange with the Mg-rich core of the pyroxene megacryst (Fe# < 0.8) to
higher amount of exchange with the olivines (all have Fe# > 0.86).
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Figure 12: Apatite and merrillite in LAR 12011 (from Howarth et al., 2016).
Figure 13: from Bellucci et al. (2017) showing variation of halogens with Cl isotopic
composition
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Figure 14 (left): Phosphate OH, F, and Cl distribution in LAR 06319 from Howarth et al. (2016).
Figure 15 (right): FMQ versus Temperature calculated for various redox equilibria in LAR
06319 (top) and La/Yb versus FMQ for select shergottites from Peslier et al (2010).
Petrology
Oxygen fugacity
Peslier et al. (2010) used olivine-spinel-orthopyroxene equilibria in LAR 06319 to
constrain the fO2 of crystallization near FMQ-2. Lower than expected from trends of
LREE enrichment and fO2 (Figure 15). They also found evidence from two oxide
oxybarometry and thermometry for late stage oxidation, perhaps linked with degassing.
Examining the Fe3+/Fe2+ ratio in maskelynite, Satake et al. (2014) found that LAR 06319
has a high Fe3+/Fe2+ ratio, higher than depleted and intermediate shergottites, and
comparable to many other enriched shergottites.
Volatiles
Much work on LAR 06319 and 12011 has focused on the phosphates and what their
detailed composition can tell us about volatiles in Mars. The controversy surrounding
phosphate chemistry includes interpretations of the role of the mantle source (Filiberto et
al., 2016; Williams et al., 2016; Sharp et al., 2016), magmatic processes (Shearer et al.,
2015), magmatic degassing (Balta et al., 2013), later hydrothermal alteration (Howarth et
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al., 2016, Bellucci et al., 2017), or shock effects (Williams et al., 2016; Sharp et al.,
2016).
Balta et al. (2013) documented a wide range of Cl contents in LAR 06319 phosphates and
argued that the wide range reflects degassing of initially water-bearing magma (Figure
16). Similarly, Filiberto et al. (2016) argue that Martian magmas were not volatile
saturated, had water/chlorine and water/fluorine ratios ~0.4–18, and are most similar, in
terms of volatiles, to terrestrial MORBs. They also argue that there are variations in the
volatile content in the Martian interior, but more bulk chemical data, especially for
fluorine and water, are required to investigate these variations.
Shearer et al. (2015) argued that the occurrence of merrillite does not imply low-volatile
component in the Martian magmas. However, the low whitlockite and bobdownsite
contents (Figure 17) suggest that these samples were not altered by hydrothermal fluids
and therefore not reset owing to aqueous fluid interactions, as argued by others.
Figure 16: Variable Cl and OH in LAR 06319 apatites from the study of Balta et al. (2013).
Table 2: Bulk compositional summary for LAR 06319 (LAR 12011 n/a)
technique
SiO2 % 46.7 (a)
TiO2 0.68 (a)
Al2O3 6 (a)
FeO 20.4 (a)
MnO 0.48 (a)
MgO 15.8 (a)
CaO 6.46 (a)
Na2O 1.14 (a)
K2O 0.14 (a)
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P2O5 0.61 (a)
S %
sum
Sc ppm 30.2 (a)
V 202 (a)
Cr 3677 (a)
Co 54.9 (a)
Ni 182 (a)
Cu 7.21 (a)
Zn 41.3 (a)
Ga 13.5 (a)
Ge ppb 1020 (a)
As
Se
Rb 5.49 (a)
Sr 44.9 (a)
Y 11.3 (a)
Zr 48.4 (a)
Nb 3.41 (a)
Mo
Ru 1.51 (b)
Rh
Pd ppb 7.4 (b)
Ag ppb
Cd ppb
In ppb
Sn ppb
Sb ppb
Te ppb
Cs ppm 0.3 (a)
Ba 24.5 (a)
La 1.73 (a)
Ce 4.25 (a)
Pr 0.57 (a)
Nd 2.94 (a)
Sm 1.1 (a)
Eu 0.47 (a)
Gd 1.55 (a)
Tb 0.3 (a)
Dy 2.01 (a)
Ho 0.42 (a)
Er 1.14 (a)
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Tm 0.16 (a)
Yb 1.15 (a)
Lu 0.15 (a)
Hf 1.28 (a)
Ta 0.12 (a)
W ppb 930 (a)
Re ppb 0.076 (b)
Os ppb 0.76 (c )
Ir ppb 0.54 (b)
Pt ppb 4.18 (b)
Au ppb
Th ppm 0.26 (a)
U ppm 0.07 (a)
References for analyses (a) ICP-MS; Sarbadhikari et al. 2009; (b) ICP-MS; Brandon et
al. 2012; (c) NTIMS; Brandon et al. 2012.
Figure 17: Variation of molar Na and Mg# for merrillites from a large suite of Martian
meteorites as well as lunar merrillites (from Shearer et al., 2015).
Howarth et al. (2016) calculate the relative volatile fugacities of the parental melts at the
time of apatite formation. Although several studies have found evidence for degassing in
the late-stage mineral assemblage of LAR 06319, the apatite evolutionary trends cannot
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be reconciled with this interpretation. The variable Cl contents and high OH contents
measured in apatites are not consistent with fractionation either. Volatile fugacity
calculations (Howarth et al., 2016) indicate that water and fluorine activities remain
relatively constant, whereas there is a large variation in the chlorine activity. They
suggest that the high and variable Cl contents and high OH contents of the apatite are the
results of post-crystallization interaction with Cl-rich, and possibly water-rich, crustal
fluids circulating in the Martian crust. Bellucci et al. (2017) showed that Cl isotope
compositions in Martian meteorites exhibit a larger range than all pristine terrestrial
igneous rocks. For example, phosphates in ancient (>4 Ga) meteorites (ALH 84001 and
NWA 7533) have positive δ37Cl anomalies (+1.1 to ‰+2.5‰); these samples also exhibit
whole rock and grain scale evidence for hydrothermal or aqueous activity. In contrast,
phosphates in the younger (<600 Ma) shergottite meteorites such as LAR 12011 have
negative δ37Cl anomalies (−0.2 to ‰−5.6‰). Phosphates with the largest negative δ37Cl
anomalies display zonation in which the rims of the grains are enriched in all halogens
and have significantly more negative δ37Cl anomalies suggestive of interaction with the
surface of Mars during the latest stages of basalt crystallization.
To shed light on the importance of degassing or fluid interaction in the crust, Udry et al.
(2016) measured Li, B, and Be abundances and Li isotope profiles in pyroxenes, olivines,
and maskelynite from four compositionally different shergottites—Shergotty, QUE
94201, LAR 06319, and Tissint—using secondary ion mass spectrometry (SIMS). All
three light lithophile elements (LLE) are incompatible: Li and B are soluble in H2O-rich
fluids, whereas Be is insoluble. Shergotty, LAR 06319, and Tissint appear to have been
affected by post-crystallization diffusion, based on diffusion modelling of their LLE and
Li isotope profiles. However, it is possible that the sub solidus diffusion overprints a
degassing pattern (Udry et al., 2016).
Clearly more work is required to distinguish between the various hypotheses to explain
the phosphate and volatile element compositional variation. The LAR 06319 and LAR
12011 samples will be crucial to the solution given their unique and strong isotopic and
compositional zoning.
Chemistry
Sarbadhikari et al. (2009) determined the bulk chemical composition and REE pattern for
LAR 06319 (Table 2, Figure 18). LAR 06319 has a composition similar to the
“enriched” shergottites, with LREE enrichment compared to the depleted shergottites
(Figure 18). Walker et al. (2009), Brandon et al. (2012), and Tait and Day (2018)
reported data on the highly siderophile element contents (Figure 19). Wang and Becker
(2017a,b) report data for Cu and Ag in LAR 06319.
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Figure 18: REE diagram for LAR 06319 (red line, stars) compared to other enriched shergottites
and olivine-phyric and lherzolitic depleted shergottites (from Sarbadhikari et al. 2009).
Figure 19: HSE contents of LAR 06319 (x) compared to other intermediate MgO shergottites
(from Brandon et al., 2012) showing the overall depletion in compatible HSE (Os, Ir, Ru) relative
to incompatible HSE (Pt, Pd, and Re).
Radiogenic age dating
Park et al. (2013) made 39Ar–40Ar (Ar–Ar) analyses of whole rock (WR) and mineral
samples of LAR 06319, in order to determine Ar–Ar age and the 40Ar/36Ar ratios of
trapped Martian Ar. Samples released trapped (excess) 40Ar and 36Ar and suggested Ar–
Ar ages older than their formation ages. Because trapped Ar components having
different 40Ar/36Ar were released at different extraction temperatures, Park et al. (2013)
utilized only a portion of the data to derive preferred Ar–Ar ages. They obtained an Ar–
Ar age 163 ± 13 Ma for LAR whole rock. They identified two trapped Ar components. At
low temperatures, particularly for plagioclase, Trapped-A with 40Ar/36Ar 285 ± 3 was
released, and they argue this is most likely absorbed terrestrial air. At high extraction
temperatures, particularly for pyroxene, Trapped-B with 40Ar/36Ar 1813 ± 127 was
released. This Ar component is Martian, and its isotopic similarity to the Martian
atmospheric composition suggests that it may represent Martian atmospheric Ar
incorporated into the shergottite melt via crustal rocks. Trapped-B partitioned into
pyroxene at a constant molar ratio of K/36ArTr = 80 ± 21 x 106 for LAR 06319. Trapped-A
mixed in different proportions with Trapped-B could give apparently intermediate
trapped 40Ar/36Ar compositions commonly observed in shergottites (see Figure 20).
Shih et al. (2009) have dated LAR 06319 by internal Rb-Sr isochron at 207 ± 14 Ma with
ISr = 0.722509 ± 69 (Figure 21) and Shafer et al. (2010) by Lu-Hf isochrons 179 ± 29 Ma
(Figure 22). Sm-Nd dating of LAR 06319 was done by Shih et al. (2009), 190 ± 29 Ma
(Figure 23), which compares well to the Sm-Nd results of Shafer et al. (2010) – 183 ± 12
Ma (Figure 24) (Table 3).
Table 3: Summary of Age Data for LAR 06319
Rb-Sr Sm-Nd Lu-Hf Ar-Ar Pb-Pb
Shih et al. (2009) 207±14 Ma 190±26 Ma
Shafer et al. (2009) 183±12 Ma 179±29 Ma
Park et al. (2013) 163±13 Ma (wr)
201±13 Ma (cpx)
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Figure 20: Ar-Ar data and isochrons for LAR 06319 whole rock (left) and pyroxene (right) from
Park et al. (2013).
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Figure 21: Shih et al. (2009) Rb-Sr isochron for LAR 06319.
Figure 22: Shafer et al. (2010) Lu-Hf isochron for LAR 06319.
Figure 23: Shih et al. (2009) Sm-Nd isochron for LAR 06319.
Figure 24: Shafer et al. (2010) Sm-Nd isochron for LAR 06319.
Cosmogenic isotopes and exposure age
Nagao and Park (2008) reported an “exposure age” of 3.3 Ma.
Stable Isotopes
Oxygen isotopes were determined by Z. Sharp and reported in Antarctic Newsletter 30
#2. Subsequent analyses reported by Sarbadhikari et al. (2009) are in agreement and both
sets of analyses are consistent with a Martian origin for LAR 06319.
Boctor et al. (2009) first reported H isotopes, including D values near 4000, using SIMS
on melt inclusion glasses from LAR 06319. Usui et al. (2012) carried out more extensive
SIMS work on olivine-hosted melt inclusions from Yamato 980459, representing a very
primitive Martian melt, and LAR 06319. They found that the Martian mantle has a
chondritic and Earth-like D/H ratio (D ~275%), with a water content of the depleted
shergottites mantle of 15–47 ppm. LAR 06319, on the other hand, exhibits an extreme
D of ~5000%, indicative of a surface reservoir (e.g., the Martian atmosphere or crustal
hydrosphere), and suggests that LAR formed by melting of an enriched mantle and later
assimilation of old Martian crust.
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Based on in situ hydrogen isotope (D/H) analyses of quenched and impact glasses in
Martian meteorites, Usui et al. (2015) provided evidence for the existence of a distinct
water/ice reservoir (D/H = ∼2–3x Earth’s ocean water) that is unlikely to be a simple
mixture of atmospheric and primordial water retained in the Martian mantle (D/H
≈Earth’s ocean water). This reservoir could instead represent hydrated crust and/or
ground ice interbedded within sediments.
Franz et al. (2008, 2014) measured S isotopes in a large set of Martian meteorites
including LAR 06319. They found it contains 1300 ppm total S, which was 750 ppm
Acid Volatile Sulfide and 550 ppm sulfate. Among most samples in their study, the
variability in 33S is similar to that in chondritic sulphur and terrestrial mantle Sulphur -
<0.01%, implying that sulphur was well mixed in the most abundant inner Solar System
materials. The 34S values derived from shergottites AVS data from their study are also
homogeneous and yielded a mean value within error of the standard (CDT) value,
suggesting efficient mixing in the Martian mantle and little or no fractionation of sulphur
isotopes during core formation. They found significant deviations from the mean in only
three shergottites, which they ascribe to photochemical Martian atmospheric effects, and
mid-crustal and near surface assimilation in some of the basaltic shergottites.
Many stable isotope studies are aimed at deciphering the building blocks of Mars, such as
Li, Ca, Mg, and Fe isotopes:
Li - LAR 06319 was used, along with a suite of Martian meteorites, to estimate the Li
isotopic composition of the Martian mantle. This work by Magna estimated the 7Li for
Mars to be 4.2 +/- 0.9, within error of the estimate for Earth’s mantle 3.5 +/- 1.0 (Magna
et al., 2015a).
Ca - Magna et al. (2015b) measured Ca isotopes in a suite of Martian meteorites,
including LAR 06319, and found that Mars has identical Ca isotopic composition to
Earth and Moon. They suggest the inner solar system is homogeneous with respect to Ca
isotopic composition.
Mg - Magna et al. (2017) examined Mg isotopic composition of a suite of Martian
meteorites and made a similar finding to that for Li and Ca – that Mars silicate mantle has
a 26Mg value that is identical to that for the rest of the inner solar system.
Fe - Sossi et al. (2016) measured Fe isotopes in a suite of Martian meteorites (including
LAR 06319) and found that when all are corrected for magmatic fractionation, the
Martian mantle has 57Fe slightly lighter than Earth, but identical to chondrites and
HEDs.
Williams et al. (2016) and Sharp et al. (2016) measured Cl isotopes in Martian meteorites
and proposed that the lightest values of 37Cl (~ -4) represent the mantle and that heavier
values up to ~ +1, represent interaction with crust or surface lithologies (Figure 25). This
interpretation is at odds with that of Bellucci et al. (2017) who measured Cl isotopes and
halogens in phosphates from Martian meteorites. They conclude that phosphates with no
textural, major element, or halogen enrichment evidence for mixing with this surface
reservoir have an average δ37Cl of ‰−0.6‰, which supports a similar initial Cl isotope
composition for Mars, the Earth, and the Moon. Oxidation and reduction of chlorine are
the only processes known to strongly fractionate Cl isotopes (both (+)ve and (-)ve),
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and perchlorate has been detected on the Martian surface, which has 5000 ppm Cl (Figure
13). Perchlorate formation and halogen cycling via brines has been active throughout
Martian history, and has a significant influence of the Cl isotopic composition (Figure
26).
Figure 25: Chlorine isotopic composition of Martian meteorites determined by Williams et al.
(2016) (left) and Sharp et al. (2016) (center). The lightest values among the Martian meteorites
are interpreted as representative of the mantle.
Figure 26 (right): Schematic illustration of the possible source of Martian light and heavy
chlorine isotopic values measured in Martian meteorites (from Bellucci et al., 2017).
Other Studies
Measurements of the short-lived isotopic systems Hf-W and Sm-Nd can be used to place
constraints on the timing of differentiation in Mars. Measurements of LAR 12011 yield
W= 0.33 ±0.10 142Nd = -0.13 ±0.06 (Kruijer et al., 2017; Figure 27). These values,
when combined with results of depleted and more enriched Martian meteorites, show that
the Martian mantle differentiated between 25 and 35 Ma after T0.
Bellucci et al. (2015, 2018) argue that the Pb isotopic composition of LAR 12011
supports derivation from a mantle with a value of 4.1-4.6, which furthermore indicates
that core formation had little influence on the mantle values. Instead the variation in
values for the Martian mantle is likely due to early mantle differentiation and sulfide
segregation (Bellucci et al., 2015, 2018).
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Figure 27: W – Nd isotopic diagram for Martian meteorites – LAR 12011 plots as a maroon
square near the “Bulk Martian mantle” values.
Processing:
Processing details of LAR 06319 and 12011 are presented here and have somewhat
parallel histories. A summary of how the samples have been subdivided, including
potted butts and thin/thick sections, chips, and remaining usable mass, is presented in
Figures 27 and 30. Both samples were initially subdivided into splits ,0, ,1, and ,3 to
prepare thin sections from a potted butt (,1) and have a chip on reserve in case oxygen
isotopic analysis is required (Figures 28 and 31). After classification and announcement
in the newsletter, the meteorites were further subdivided to fill sample requests. Images
of typical subsplits of LAR 06319 and LAR 12011 are presented in Figures 29 and 32).
For LAR 06319, 70% of original mass, and 22 thin sections remain for future studies,
while for LAR 12011, 97% of original mass, and 14 thin sections remain for future
studies.
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Figure 28: Chart showing a summary of the subdivision of LAR 06319, including initial
processing and later subdivision for allocations to PIs. Remaining 55.12 g of material represents
70% of original mass.
Figure 29: Image of LAR 06319 after initial processing and the splits 0, 1, and 3.
Figure 30: Typical subsplits photos of LAR 06319 showing splits 16-22, and ,0 (left) and splits 7
through 11, selected from ,3 (right).
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Figure 31: Chart showing a summary of the subdivision of LAR 12011, including initial
processing and later subdivision for allocations to PIs. Remaining 678.65 g of material
represents 97% of original mass.
Figure 32: Images of LAR 12011 after initial processing and the splits 0, 1, and 3.
Figure 33: Typical subsplits photos of LAR 12011 showing splits 29 and 30 derived from 3 (left)
and splits 24 and 25 derived from 3 (right).
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References for LAR 06319
Balta, J. B., Sanborn, M., Mcsween, H. Y., Wadhwa, M. (2013) Magmatic history and
parental melt composition of olivine-phyric shergottite LAR 06319: Importance of
magmatic degassing and olivine antecrysts in Martian magmatism. Meteoritics &
Planetary Science 48, 1359-1382, http://dx.doi.org/10.1111/maps.12140.
Bellucci, J. J., Nemchin, A. A., Whitehouse, M. J., Snape, J. F., Bland, P., & Benedix, G.
K. (2015) The Pb isotopic evolution of the Martian mantle constrained by initial Pb in
Martian meteorites. Journal of Geophysical Research: Planets 120, 2224-2240.
Bellucci, J. J., Nemchin, A. A., Whitehouse, M. J., Snape, J. F., Bland, P., Benedix, G.
K., & Roszjar, J. (2018) Pb evolution in the Martian mantle. Earth and Planetary
Science Letters 485, 79-87.
Bellucci, J. J., Whitehouse, M. J., John, T., Nemchin, A. A., Snape, J. F., Bland, P. A., &
Benedix, G. K. (2017) Halogen and Cl isotopic systematics in Martian phosphates:
Implications for the Cl cycle and surface halogen reservoirs on Mars. Earth and
Planetary Science Letters 458, 192-202, https://doi.org/10.1016/j.epsl.2016.10.028.
Boctor N.Z., Wang J., Alexander C.M.O., Steele A. and Armstrong J. (2009) Hydrogen
isotope signatures and water abundances in nominally anhydrous minerals from the
olivine-phyric Shergottite LAR 06319 (abs#5176). Meteorit. & Planet. Sci. 44, A35.
Brandon, A. D., Puchtel, I. S., Walker, R. J., Day, J. M. D., Irving, A. J., Taylor, L. A.
(2012) Evolution of the Martian mantle inferred from the 187Re-187Os isotope and
highly siderophile element abundance systematics of shergottite meteorites,
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