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Icarus 175 (2005) 32–35www.elsevier.com/locate/icaru

Isotopic fractionation of methane in the martian atmosphere

Hari Naira,∗, Michael E. Summersb, Charles E. Millerc, Yuk L. Yungd

a Mail Stop 168-414, Instruments and Science Data Systems Division, Jet Propulsion Laboratory, Pasadena, CA 91109, USAb School of Computational Sciences and Department of Physics and Astronomy, George Mason University, Fairfax, VA 22030, USA

c Earth and Space Sciences Division, Jet Propulsion Laboratory, Pasadena, CA 91109, USAd Division of Geological and Planetary Sciences, California Institute of Technology, Pasadena, CA 91125, USA

Received 17 August 2004; revised 26 October 2004

Available online 21 January 2005

Abstract

The existence of methane in the martian atmosphere may be an indicator of subsurface life. Biological processes are known tothe common isotopologues of methane, and hence measuring these isotopic ratios may yield constraints on the nature of the methAny measurement of the isotopic ratios of atmospheric methane must consider the additional fractionation due to photochemistryquantify the isotopic ratios of the source. Using a one-dimensional photochemical model, we find that photochemistry has a sma.5�)contribution toδ13C(CH4) but has a large (114�) contribution toδD(CH4). Confirmation of these fractionation values will require additiolaboratory data on key model inputs, particularly the ultraviolet absorption cross sections of13CH4 and kinetic rate coefficients for threactions of13CH4 and CH3D with OH and O(1D) at pressures and temperatures relevant to the martian atmosphere. 2004 Elsevier Inc. All rights reserved.

Keywords: Exobiology; Mars, atmosphere; Photochemistry

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1. Introduction

The recent tentative detection of methane in the mtian atmosphere(Formisano et al., 2004; Krasnopolskyal., 2004; Mumma et al., 2003)raises the question of itorigin, as the martian atmosphere is an oxidizing envirment where the lifetime of methane is on the order200 years. A number of possible sources have beengested, such as subsurface biology(Weiss et al., 2000Summers et al., 2002), volcanism(Wong et al., 2003), orcometary impacts(Kress and McKay, 2004).

It is possible to discriminate between terrestrial methsources by measuring the relative abundance of the tcommon isotopologues CH4, CH3D, and 13CH4. For ex-ample, biogenic methane generally hasδ13C(CH4) lighterthan−60� andδD(CH4) lighter than−150�, while ther-mogenic (e.g., volcanic) methane generally hasδ13C(CH4)

* Corresponding author. Fax: +1-818-393-4802.E-mail address: hari.nair@jpl.nasa.gov(H. Nair).

0019-1035/$ – see front matter 2004 Elsevier Inc. All rights reserved.doi:10.1016/j.icarus.2004.10.018

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heavier than−50� and δD(CH4) heavier than−275�(Cicerone and Oremland, 1988; Tyler, 1992). The referenceisotopic standards forδ13C andδD are Pee Dee belemniand Standard Mean Ocean Water, respectively. Here weused theδ symbol to represent the isotopic enhancemena sample, defined by

(1)δ = 1000[(r/R) − 1

],

wherer is the isotopic ratio (e.g., D/H) of the sample aR is the isotopic ratio of the reference. The units ofδ are“permil” (�), or parts per thousand.

Photochemistry will additionally fractionate atmosphemethane. Based on the source and atmospheric vof terrestrial δD(CH4), Tyler (1992) inferred an isotopicshift for δD(CH4) of 210� due to atmospheric chemistrSaueressig et al. (2001)found enhancements of up to 30�for δ13C(CH4) and 300� for δD(CH4) in their photochem-ical model of the terrestrial stratosphere, in good agreem

with the observed isotopic ratios.

If the degree of photochemical fractionation of themethane family can be quantified, measurements of these

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isotopic ratios in the martian atmosphere can provide astraint on the isotopic ratios of the source, and possiblytinguish biological sources from non-biological ones(Yungand DeMore, 1999; Summers et al., 2002).

2. Photochemical model

Our one-dimensional photochemical model is identicacase “f” of Nair et al. (1994), with the addition of CH4,CH3D, and13CH4. Surface fluxes for the methane specare specified as boundary conditions. We assume thatare no photochemical sources for these species and thonly losses are due to photolysis or reaction with O(1D) orOH.

Photochemical cross sections for CH4 and CH3D aretaken fromLee et al. (2001). Since the photochemical crosections for13CH4 have not been measured, we estimthem using the differences in the ground state zero pvibrational energies (ZPEs) between CH4 and 13CH4. Weassume a unit quantum yield for photodissociation asultraviolet photon energies are significantly larger thanC–H bond strength. This method has been previously uto estimate photochemical cross sections of the heaviertopologues of N2O (Yung and Miller, 1997).

The ZPEs for CH3D and 13CH4 relative to CH4 are−599± 1 and−25.5 ± 0.5 cm−1, respectively. These corespond to blue shifts of 0.9 and 0.04 nm, respectively120 nm. The CH3D cross sections predicted by this methagree well with the measured values ofLee et al. (2001)from105 to about 130 nm, but are smaller at longer wavelenwhere the cross sections are changing rapidly. The ZPEfor 13CH4 is small enough that the predicted cross sectido not significantly differ from those of CH4. The uncer-tainty in the13CH4 cross sections is thus dominated byassumption that the CH3D and 13CH4 cross sections asfunction of wavelength have identical shapes.

Rate coefficients for the reactions of CH3D and 13CH4with O(1D) and OH were computed using the kinetic isotoeffect (KIE) factors measured bySaueressig et al. (2001)andthe corresponding CH4 reaction coefficients fromDeMore

Table 1Rate coefficients for methane loss reactions

Reaction Rate coefficient

CH4 + hν 2.924× 10−6

CH4 + O(1D) 1.5× 10−10

CH4 + OH 2.45×10−12e−1775/T

CH3D + hν 2.908× 10−10

CH3D + O(1D) 1.42× 10−10

CH3D + OH 2.23×10−12e−1824/T

13CH4 + hν 2.924× 10−6

13CH4 + O(1D) 1.48× 10−10

13CH4 + OH 2.45×10−12e−1775/T

The units for photolysis and two-body reactions are s−1 and cm3 s−1, re-spectively.

tion on Mars 33

ee

et al. (1997). Table 1shows these rate coefficients, as was the photodissociation coefficient calculated at the tothe model atmosphere (240 km). At 200 K, the uncertain the rate coefficients for the reaction of CH4 with O(1D)and OH are 40 and 30%, respectively(DeMore et al., 1997).The 2σ uncertainties in the KIE values are on the order ofew percent(Saueressig et al., 2001). The KIE uncertaintiesare more important for this study, as we are interested inrelative loss rates among the methane family memberstheir absolute rates.

Surface fluxes specified as lower boundary conditionseach methane species are listed inTable 2. The flux for CH4(φ0) was chosen to reproduce the observed surface miratio of 10−8 (Formisano et al., 2004; Krasnopolsky et a2004). The fluxes for CH3D (φ1) and13CH4 (φ2) were de-termined by

φ1 = 4∗ fD1 ∗ fD2 ∗ fD3 ∗ φ0,

φ2 = fC1 ∗ fC2 ∗ fC3 ∗ φ0,

wherefD1 = 1.5576× 10−4 is the terrestrial D/H standarvalue(Fegley, 1995), fD2 = 5.2 is the enrichment factor fodeuterium on Mars relative to Earth(Bjoraker et al., 1989Owen et al., 1988; Krasnopolsky et al., 1997), fD3 = 0.710is the enrichment factor for a biogenic source(Tyler, 1992),fC1 = 1.1235× 10−2 is the terrestrial13C/12C standardvalue (Fegley, 1995), fC2 = 1 is the enrichment factor fo13C on Mars relative to Earth(Nier and McElroy, 1977), andfC3 = 0.943 is the enrichment factor for a biogenic sou(Tyler, 1992). There is significant (up to 100�) uncertaintyin the value of martian atmosphericδ13C (Jakosky, 1991),but for the purposes of this paper a precise value isrequired. The enrichment factors for biogenic sources arsumed to be the same as the terrestrial values.

3. Results

The calculated surface mixing ratios, column abdances, and column-integrated lifetimes for the methspecies are shown inTable 3. The column-integrated losrates and timescales for each photochemical sink are shin Table 4. Profiles of all other species in the model anearly unchanged from those in case “f” ofNair et al. (1994),as the addition of trace amounts of methane is an insigcant perturbation to the standard model.

Since photolysis is not effective as a destruction mecnism below∼ 60 km, we find a lifetime of 380 years at thsurface. This is longer than the surface lifetime of 300 ye

Table 2Lower boundary conditions

Species Surface flux (cm−2 s−1)

CH4 3.460×105

CH3D 7.960×102

13CH4 3.670×103

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found bySummers et al. (2002), who used a rate coefficienfor the reaction O+ HO2 that is 50% of that used byNair etal. (1994). When using this value, we obtain the same sface lifetime of 300 years. O+ HO2 is the dominant sinkfor odd oxygen (O and O3) in the martian atmosphere. Uing a smaller rate coefficient for this reaction increasesamount of atmospheric O3 and hence increases the rateCH4 destruction by O(1D), which is produced by photolysiof O3.

From the column-integrated lifetimes, we see that ptolysis is the largest contributor to methane destruction,does not discriminate between the different isotopologReaction with OH is the second most important destrucpathway, and removes CH4 and 13CH4 significantly fasterthan CH3D. Reaction with O(1D) is marginally faster forCH4 and13CH4 than for CH3D.

In order to computeδ values using Eq.(1), we taker

to be the ratio of concentrations (e.g., [CH3D]/[CH4]) andR to be the ratio of surface fluxes (e.g.,φ1/φ0). From thefluxes and calculated mixing ratios inTables 2 and 3, wefind thatδD(CH4) = 114� andδ13C(CH4) = 4.5� at thesurface. These values remain nearly constant up to 100as shown inFig. 1. Above∼ 120 km,13CH4 and CH3D falloff in abundance relative to the lighter CH4 due to moleculardiffusion. We have not included reactions of methane wionic species in this model, which may affect the isotofractionation in the ionosphere and above.

4. Conclusions

If methane is present in the martian atmosphere, a msurement of the relative abundances of its common isotlogues13CH4 and CH3D may serve to constrain the natuof its source. These atmospheric measurements will habe corrected for photochemical processes which will mothe relative abundances of the methane species fromsource values.

Table 3Surface mixing ratios, column amounts, and column lifetimes

Species Surface mixing ratio Column abundance(cm−2)

τ

(years)

CH4 9.993× 10−9 2.305× 1015 211

−11 12

CH3D 4.347× 102 431 2.010×102

13CH4 1.808× 103 430 1.165×103

75 (2005) 32–35

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We expect the lower atmosphere to be enriched inheavier methane isotopologues relative to the source.measured atmospheric13CH4/CH4 ratio should be veryclose(4.5�) to the emissions by the source, while the msured atmospheric CH3D/CH4 ratio should be appreciablhigher(114�) than the source emissions, primarily duethe more rapid reaction of CH4 with OH.

Figure 2 shows how the measured isotopic ratioatmospheric methane should differ from the isotopictio of the source. We take the source composition toδ13C(CH4) = −57� andδD(CH4) = −290�, the same asfor terrestrial sources(Tyler, 1992). In this case the measuredδ13C(CH4) would be−52� andδD(CH4) would be−176�.

Fig. 1. Calculated vertical profiles ofδ13C(CH4) andδD(CH4).

as

CH3D 2.561× 10 5.907× 10 23513CH4 1.065× 10−10 2.456× 1013 212

Fig. 2. Isotopic composition of a CH4 source (assumed to be the sameterrestrial) and the expected atmospheric composition.

Table 4Column integrated loss rates and timescales for methane isotopologues

hν column rate(cm−2 s−1)

τ

(years)OH column rate(cm−2 s−1)

τ

(years)O(1D) column rate(cm−2 s−1)

τ

(years)

CH4 1.700× 105 430 1.098×105 665 6.702× 104 1090

931 1.620× 102 1155668 7.048× 102 1104

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Significant uncertainties in this study are the photolyrate of13CH4, as well as values for the kinetic isotope effefor the reactions of13CH4 and CH3D with OH and O(1D).Further laboratory measurements will be extremely usin order to design and conduct experiments that can meathe isotopic fractionation of CH4 on Mars to the requiredprecision to detect biogenic signatures.

Acknowledgments

We thank M. Gerstell for a critical reading of the manscript. This work was supported in part by an NSF grantthe NASA Astrobiology Institute at JPL.

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