1

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-

2

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.

3

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

4

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).

5

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).

6

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).

7

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

8

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

9

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)

10

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)

11

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

12

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)

14

Figure 20: Ar-Ar data and isochrons for LAR 06319 whole rock (left) and pyroxene (right) from

Park et al. (2013).

15

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.

16

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),

17

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).

18

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.

19

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).

20

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).

21

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,

Geochimica et Cosmochimica Acta 76, 206-235;

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