 PRELIMINARY INTERPRETATIONS OF ATMOSPHERIC STABLE ISOTOPES AND ARGON FROM MARS SCIENCE LABORATORY (SAM).  J.H. Jones1, P.B. Niles1, C.R. Webster2, P.R. Mahaffy3, G.J. Flesch2, L.E. Christensen2, L.A. Leshin4, H. Franz3, M. Wong5, S.K. Atreya5, P.G. Conrad3, H. Manning6, R. Navarro-Gonzalez7, T. Owen8, B. Pepin9, J.C. Stern3, M. Trainer3, S.P. Schwenzer10 and the MSL Science Team, 1NASA/JSC, Houston, TX 77058 (john.h.jones@nasa.gov), 2Jet Propulsion Laboratory, Caltech, 4800 Oak Grove Dr., Pasadena, CA 91109, 3NASA Goddard Space Flight Center (GSFC), Greenbelt, MD 20771, 4Rensselaer Polytechnic Institute, Troy, NY 12180, 5Univ. of Michigan, MI 48105, 6Concordia College, Moorhead, MN, 7Universidad Nacional Autónoma de México, Mexico City, Mexico, 8Univ. of Hawaii, Honolulu, HI 96822, 9Univ. of Minnesota, Minneapolis, MN, 10The Open University, Milton Keynes, UK.  Introduction:  The first analysis of the martian atmosphere was performed by the Viking missions [1, 2]. The elevated 129Xe/132Xe ratio measured by Viking allowed [3] to note the similarity of the Xe extracted from quenched shock melts in the EETA 79001 Antarctic shergottite with the Viking measurements. This was the first solid evidence that meteorites had been delivered from Mars to the Earth [3, 4].  The timing of this discovery was serendipitous. Only two years before, a lunar meteorite had been found in Antarctica. Thus, there was now solid evidence that impacts on other planets could deliver material to the Earth. In addition to the large 129Xe anomaly, Viking also noted that the 40Ar/36Ar of the martian atmosphere was ~3000 and that δ15N was about +600‰  [1, 2]. Earthbased spectroscopic measurements subsequently indicated that the D/H ratio of the martian atmosphere (δD) was ~+4000‰ [5]. The combination of isotopically heavy N and H led to speculation that the isotopic composition of these elements had been modified by atmospheric escape processes [6]. The chemical and isotopic composition of the martian atmosphere was further refined by laboratory measurements of martian meteorites.  One example of this is the 40Ar/36Ar of the martian atmosphere, where more refined analyses of EET 79001 shock glasses yielded a 40Ar/36Ar ratio of ~1900, as opposed to the Viking measurement of ~3000 [7, 8]. Another example is that of 36Ar/38Ar = <4 in these glasses, which is significantly lower than the chondritic ratio of ~5.4. This also seems to support the idea of preferential atmospheric loss of lighter Ar isotopes over geologic time [7-9].  However, other measurements of CO2 isotopes have not been consistent with this atmospheric loss story. Viking measured δ13C and δ18O values of 0±	  50‰. Earth-based spectroscopy has in fact suggested depleted values for δ13C of -22±	  20‰ [10]. Finally, the recent Phoenix lander measured a δ13C for CO2 in the martian atmosphere of -2‰ [11]. This is in contrast to measurements of trapped gas in martian meteorite EETA 79001 which yielded a δ13C of +36 ± 10‰ [12].  Thus, a combination of robotic, laboratory, and telescopic measurements have set the stage for the SAM/MSL analysis of the martian atmosphere. The SAM analyses are performed in two very different ways within a single instrument package:  (i) Tunable Laser (TLS) measurements of 18O/16O and 13C/12C in atmospheric CO2; and (ii) Quadrupole mass spectroscopy (QMS) of  13C/12C, and 40Ar/36Ar.  TLS has also measured δD both in the Martian atmosphere and as evolved from a heated soil sample from Rocknest [13], with results currently under refinement. Preliminary Results:  TLS results yield an isotopic composition of atmospheric CO2 to be: +48±6‰ in δ18O and +45±4‰ for δ13C [14]. QMS results for δ13C are also in agreement with an average of +40.4±15.5‰	   [15]. These measurements agree well with previous δ18O results and are similar to previous martian meteorite measurements of δ13C. By QMS the 40Ar/36Ar ratio of the atmospere is ~1900, in good agreement with martian meteorite studies [7, 8].  Refinement of this measurement is also underway, but it is clear that the martian atmospheric Ar has been highly influenced by degassing of radiogenic 40Ar from the crust. δD measurements by TLS on water in the atmosphere yielded values of +5000 to +7000‰ [14], in good agreement with both telescopic measurements and with a new ion probe analysis of D/H in an olivine-hosted melt inclusion in the shergottite LAR 06319 [9]. In addition δ18O of atmospheric water was measured at +50 ±10‰ by the TLS [14].  Implications:  Taken at face value, our measured  enrichments in the heavy isotopes of O, C, and H supports the paradigm of atmospheric loss processes at Mars. Estimates of thermal and non-thermal escape processes during Mars' initial history suggest that early loss was catastrophic, suggesting that very little atmosphere survived earliest accretion and outgassing [9, 16]. Minerals in ALH 84001 whose crystallization and alteration ages have been dated to be near 4 Ga contain our best record of these events and contain enriched C and H isotopes. Measurements of ALH 84001 minerals show δ13C ~+40‰, and δD values ~+3000‰ [17-19]. In this view, the δ13C and δD of the martian atmosphere has not changed much over about 4 Ga. This would also be consistent with the measurement of a large enrichment of δ18O in atmospheric water vapor which could suggest that water on Mars is not in equilibrium with the crust and has also been enriched in heavy isotopes through atmospheric loss. An alternate interpretation of the data suggests that buffering of surface volatiles by crustal reservoirs could play an important role in the evolution of the atmosphere as well as its isotopic composition.  Crustal reservoirs may represent a much larger volume of volatiles whose isotopic composition was largely established very early. Thus, exchanges between atmosphere and crustal reservoirs could provide a viable means for preserving the fingerprint of ancient processes observed in both the ALH 84001 meteorite and in the modern atmosphere. The δ18O of martian CO2 is very heavy with respect to silicates in martian meteorites but is remarkably similar to terrestrial CO2.  Thus, this 18O enrichment could be a signature either of low temperature equilibration between water and CO2 [11] or enrichment through atmospheric loss. The oxygen isotopic composition of CO2 in the martian atmosphere and in carbonates from martian meteorites indicate that some sort of buffering has indeed taken place. Jakosky and Jones [20] argued, based on the Viking measurement of δ18O ~ 0‰, that oxygen could be buffered by hydrothermal interaction between the crust and atmosphere. The discovery of extensive clay mineral formation in the ancient crust of Mars [21] is evidence for substantial ancient exchange of oxygen isotopes between water and the silicate crust. This buffering between crust and water is not indicated by our water vapor δ18O data.  However, this might be explained by atmospheric exchange between CO2 and H2O with the high δ18O values being transferred from the more abundant 18O-rich CO2 to the less abundant water vapor.  While the current reservoir of CO2 on Mars is small, ~12 mbar [22], substantial CO2 may be stored as carbonate in crustal materials and may be as large as 13 bars [23]. The size and isotopic composition of this reservoir remains a large unknown. Some younger martian meteorites have at least trace levels of carbonates with low δ13C that have been a interpreted to be martian [24-26], suggesting that modern carbonate formation is ongoing [11]. It is expected that volcanic degassing and carbonate precipitation should act to decrease the carbon isotopic composition of atmospheric CO2, but exchange with an ancient, high-δ13C crustal reservoir may provide a means for explaining why δ13C has remained heavy through time.  In contrast, the low 36Ar/38Ar from martian meteorite gas probably requires loss over geologic time.  If the escape modeling is correct, minor gases like N, Ne, and Ar were totally removed following atmospheric collapse near 4 Ga and then replaced by later outgassing of the martian interior.  It is expected that degassed 36Ar/38Ar was initially chondritic/solar and then that ratio became sub-chondritic as 36Ar was preferentially lost [6].  The enriched 15N signature of the atmosphere is also attributed to this process. Summary: Given the broad agreement between C, H, and O isotopic ratios in the modern atmosphere and the ALH 84001 meteorite, it is possible that these reservoirs were established after early atmospheric loss prior to 4 Ga. The preservation of these signals over this long period of history can be explained in several slightly different ways: 1) C, O, and H have remained static in the atmosphere and have not exchanged with the surface over the past 4 Ga; 2) C, O, and H in the atmosphere have potentially varied widely over history but have been continually buffered by larger reservoirs in the crust which have remained unchanged over the past 4 Ga. This second possiblility allows for potentially large variations in atmospheric pressure to occur as CO2 is recycled back into the atmosphere from crustal reservoirs or degassed from the mantle. Future Work:  A planned SAM experiment that concentrates minor atmospheric constitutents may allow the measurement of the isotopic compositions of N2, 38Ar, Kr and Xe. Measurement of those isotopes that are most easily influenced by spallation and ncapture would allow determination of the influence of impact vaporization processes on the martian atmosphere.  In particular, the abundances of 80Kr, 82Kr, 124Xe, and 126Xe should be especially sensitive to cosmogenic production in the martian crust. References: 1. McElroy M.B., et al. (1976) Science, 194, 1295-1298. 2. Nier A.O., et al. (1976) Science, 194, 68-70. 3. Bogard D.D. and Johnson P. (1983) Science, 221, 651-654. 4. Pepin R.O. (1985) Nature, 317, 473-475. 5. Krasnopolsky V.A., et al. (1998) Science, 280, 1576-1580. 6. McElroy M.B., et al. (1976) Science, 194, 70-72. 7. Garrison D.H. and Bogard D.D. (1998) Meteoritics & Planetary Science, 33, 721-736. 8. Garrison D. and Bogard D. (2000) Meteoritics and Planetary Science Supplement, 35, 58. 9. Pepin R.O. (1994) Icarus, 111, 289-304. 10. Krasnopolsky V.A., et al. (2007) Icarus, 192, 396-403. 11. Niles P.B., et al. (2010) Science, 329, 1334-1337. 12. Carr R.H., et al. (1985) Nature, 314, 248-250. 13. Leshin L., et al. (2013) LPSC. 14. Webster C.R. et al. (2013) LPSC. 15. Franz H., et al. (2013) LPSC. 16. Lammer H., et al. (2012) Space Science Reviews, 1-42. 17. Leshin L.A., et al. (1996) Geochimica et Cosmochimica Acta, 60, 2635-2650. 18. Greenwood J.P., et al. (2008) Geophysical Research Letters, 35, 05203. 19. Romanek C.S., et al. (1994) Nature, 372, 655-657. 20. Jakosky B.M. and Jones J.H. (1994) Nature, 370, 328-329. 21. Bibring J.P., et al. (2006) Science, 312, 400-404. 22. Phillips R.J., et al. (2011) Science. 23. Bandfield J.L., et al. (2003) Science, 301, 1084-1087. 24. Clayton R.N. and Mayeda T.K. (1988) Geochimica et Cosmochimica Acta, 52, 925-927. 25. Jull A.J.T., et al. (1997) Journal of Geophysical ResearchPlanets, 102, 1663-1669. 26. Wright I.P., et al. (1992) Geochimica Et Cosmochimica Acta, 56, 817-826.  
