Using Biogenic Sulfur Gases as Remotely Detectable Biosignatures on Anoxic Planets

Shawn D Domagal-Goldman; Victoria S Meadows; Mark W Claire; James F Kasting

doi:10.1089/ast.2010.0509

. 2011 Jun;11(5):419–441. doi: 10.1089/ast.2010.0509

Using Biogenic Sulfur Gases as Remotely Detectable Biosignatures on Anoxic Planets

Shawn D Domagal-Goldman ^1,^✉, Victoria S Meadows ^2,,³, Mark W Claire ^2,,³, James F Kasting ^3,,⁴

PMCID: PMC3133782 PMID: 21663401

Abstract

We used one-dimensional photochemical and radiative transfer models to study the potential of organic sulfur compounds (CS₂, OCS, CH₃SH, CH₃SCH₃, and CH₃S₂CH₃) to act as remotely detectable biosignatures in anoxic exoplanetary atmospheres. Concentrations of organic sulfur gases were predicted for various biogenic sulfur fluxes into anoxic atmospheres and were found to increase with decreasing UV fluxes. Dimethyl sulfide (CH₃SCH₃, or DMS) and dimethyl disulfide (CH₃S₂CH₃, or DMDS) concentrations could increase to remotely detectable levels, but only in cases of extremely low UV fluxes, which may occur in the habitable zone of an inactive M dwarf. The most detectable feature of organic sulfur gases is an indirect one that results from an increase in ethane (C₂H₆) over that which would be predicted based on the planet's methane (CH₄) concentration. Thus, a characterization mission could detect these organic sulfur gases—and therefore the life that produces them—if it could sufficiently quantify the ethane and methane in the exoplanet's atmosphere. Key Words: Exoplanets—Biosignatures—Anoxic atmospheres—Planetary atmospheres—Remote life detection—Photochemistry. Astrobiology 11, 419–441.

1. Introduction

The search for life may soon expand beyond the boundaries of our solar system via the detection of spectral features of “biosignature” gases on extrasolar planets (European Space Agency, 2010; Jet Propulsion Laboratory, 2010; New Worlds Observer Team, 2010). For a gas to be a biosignature it must have a biological production rate that far outpaces abiotic sources and an atmospheric lifetime that allows it to build up to detectable levels. To be detectable, the biosignature gas must have spectral features that are (1) within a wavelength region that can be covered by instrumentation, (2) larger than the signal-to-noise ratio (S/N) for these instruments, and (3) distinguishable from other spectral features.

For biospheres in which primary productivity is dominated by oxygenic photosynthesis (henceforth referred to as “oxic” biospheres), a number of gases have been identified that meet these criteria: oxygen (O₂), ozone (O₃), or both in the presence of reduced species such as methane (CH₄) (Lovelock, 1965; Des Marais et al., 2002); nitrous oxide (N₂O) (Sagan et al., 1993); and methyl chloride (CH₃Cl) (Segura et al., 2005). The latter two gases are more difficult to detect in Earth's present atmosphere than are the first two; however, they might be more visible in the atmospheres of oxic Earth-like planets orbiting M stars due to longer atmospheric lifetimes resulting from lower photolysis rates (Segura et al., 2005).

Other biosignatures are needed for detection of “anoxic biospheres” that harbor life but not detectable amounts of atmospheric O₂ and O₃. Biogenic CH₄ could be abundant enough to be detectable in such an atmosphere (Kasting et al., 1983, 2001; Kasting, 2005; Kharecha et al., 2005; Kaltenegger et al., 2007), but its interpretation would be ambiguous because abiotic processes such as serpentinization can also produce CH₄ (Berndt et al., 1996; Kasting and Catling, 2003).

From the early history of life on Earth, we know that anoxic biospheres are possible. Studies of early Earth suggest that life was present well before significant O₂ accumulated in the atmosphere (Schopf, 1983; Holland, 1984; Farquhar and Wing, 2003; Westall, 2005; Farquhar et al., 2007). This period had vigorous biological activity without significant O₂ buildup and may have lasted as long as 1.5 billion years, approximately one-third of Earth's history. This suggests that planets with life, but without O₂/O₃, could represent a large fraction of inhabited planets. Thus, the absence of O₂/O₃ should not be taken as evidence that life does not exist on a planet's surface.

Furthermore, some planets and biospheres will not exhibit the more general feature of photochemical disequilibrium previously proposed as a universal biosignature (Lederberg, 1965; Lovelock, 1965; Des Marais et al., 2002). Unlike Earth's modern-day ecosystem, global anoxic ecosystems may drive an atmosphere toward equilibrium. For example, in the anoxic Archean biospheres considered by Kharecha et al. (2005), methanogens and acetogens combine H₂ and CO with CO₂ and H₂O to produce CH₄. They can make a metabolic living by doing this because CH₄ has a lower Gibbs free energy and hence is thermodynamically stable in such a system. The biogenic gases released from such a biosphere result from a drive toward equilibrium, not disequilibrium. Because cases like these could complicate interpretation, it is important to identify additional biosignature gases that might be signs of anoxic biospheres. In this paper, we test the ability of various gases with carbon-sulfur bonds to act as remotely detectable biosignatures for anoxic, inhabited surface environments.

The biosignature potential of S-bearing gases was reviewed by Pilcher (2003), who focused on gases with bonds between methyl groups (−CH₃) and sulfur: methanethiol (CH₃SH, also known as methyl mercaptan), dimethyl sulfide (CH₃SCH₃ or DMS), and dimethyl disulfide (CH₃S₂CH₃ or DMDS). More recently, Vance et al. (2011) suggested that CH₃SH could be used as an in situ signature for life on Mars. On modern Earth, the production of these species is dominated by biota, but they are rapidly destroyed by photolysis and by reaction with hydroxyl (OH) radicals (Kettle et al., 2001), and do not build up to concentrations detectable across interstellar distances. In this work, we consider these gases, along with carbon disulfide (CS₂) and carbonyl sulfide (OCS, sometimes abbreviated in other work as COS), two other biogenic gases that contain carbon-sulfur bonds. These two species—particularly OCS—also have volcanic and photochemical sources, but they are far smaller than biological fluxes. We henceforth use the term “S_org” as shorthand to refer to the entire suite of biologically produced species with carbon-sulfur bonds (DMS, DMDS, CH₃SH, CS₂, and OCS). Although hydrogen sulfide (H₂S) is another S-bearing gas produced by biota, large quantities of this species enter the atmosphere via volcanism. Thus, we do not consider it here as a biosignature. However, we do consider the possibility that volcanic H₂S could act as a “false positive” for biogenic S_org, as this abiotic H₂S could react in the atmosphere to form S_org species. Other work has explored the spectral signatures of sulfur dioxide (SO₂) and H₂S in detail (Kaltenegger and Sasselov, 2010), so we limit our discussion to their potential to be false positives for biological S_org production. No study to date has predicted the concentrations of all the S_org species in an anoxic atmosphere, nor has any study predicted the spectral features associated with these gases. We used a photochemical model to calculate vertical profiles of these gases for a variety of astronomical and biological contexts, and used a radiative transfer model to predict the spectral features consistent with those profiles.

2. Methods

2.1. Photochemical code

We modified the one-dimensional (altitude), low-O₂ photochemical code originally developed by Kasting et al. (1979) to study the anoxic early Earth. The numerics of this model are described by Kasting and Ackerman (1985), and the chemistry was most recently modified by Pavlov et al. (2001). We have updated this code, adding seven long-lived chemical species that have lifetimes longer than the time scale for vertical mixing: CH₃SH, DMS, DMDS, OCS, CS₂, methylthiol (CH₃S), and carbon monosulfide (CS). We also added three short-lived species, which are solved in photochemical equilibrium without considering vertical transport: excited-state CS₂, OCS₂, and HCS. These 10 species were incorporated into the chemical scheme by adding 73 chemical reactions. The current model contains 83 chemical species, 46 of which are long lived, connected by 433 chemical reactions. Additionally, many of the 360 reactions from prior work have updated reaction rate constants. A complete list of model reactions, reaction rate constants, and references can be found in Table 1.

Table 1.

List of Reactions in the Photochemical Code, along with the Reaction Rate Constants Used and a Source for the Reaction Rate Constant

Rxn. #	Reaction	Reaction rate constant	Reference
1	OCS+CH→CO+HCS	1.99·10⁻¹⁰×e^−190/T	Zabarnick et al.,1989
2	OCS+H→CO+HS	9.07·10⁻¹²×e^−1940/T	Lee et al.,1977
3	OCS+O→S+CO₂	8.3·10⁻¹¹×e^−5530/T	Singleton and Cvetanovic, 1988
4	OCS+O→SO+CO	2.1·10⁻¹¹×e^−2200/T	Toon et al.,1987
5	OCS+OH→CO₂+HS	1.1·10⁻¹³×e^−1200/T	Atkinson et al.,2004
6	OCS+OH→HS+CO₂	1.1·10⁻¹³×e^−1200/T	Atkinson et al.,2004
7	OCS+S→CO+S₂	1.5·10⁻¹⁰×e^−1830/T	Schofield, 1973
8	OCS+S+M→OCS₂+M	8.3·10⁻³³×den	Basco and Pearson, 1967
9	OCS₂+CO→OCS+OCS	3.0·10⁻¹²	Zahnle et al.,2006
10	OCS₂+S→OCS+S₂	2.0·10⁻¹¹	Zahnle et al.,2006
11	C₂H₆S+CH₃→CH₄+C₂H₄+HS	6.92·10⁻¹³×e^−4610/T	Arthur and Lee, 1976
12	C₂H₆S+H→C₂H₅+H₂S	8.49·10⁻¹²×e^−1200/T	Lam et al.,1989
13	C₂H₆S+H→CH₃SH+CH₃	4.81·10⁻¹²×e^−1100/T×(T/300)^1.7	Zhang et al.,2005
14	C₂H₆S+H→H₂+C₂H₄+HS	8.34·10⁻¹²×e^−2212/T×(T/300)^1.6	Zhang et al.,2005
15	C₂H₆S+OH→H₂O+C₂H₄+HS	1.13·10⁻¹¹×e^−253/T	Atkinson et al.,2004
16	C₂H₆S₂+H→CH₃SH+CH₃S	9.47·10⁻¹²×e^−50/T	Ekwenchi et al.,1980
17	CH₃+HS→CH₃SH	1.66·10⁻¹¹	Shum and Benson, 1985
18	CH₃S+CH₃S→C₂H₆S₂	4.00·10⁻¹¹	Anastasi et al.,1991
19	CH₃S+CO→CH₃+OCS	2.6·10⁻¹¹×e^−5940/T	Assumed same as k(CH₃O+CO)
20	CH₃S+CS→CH₃+CS₂	2.6·10⁻¹¹×e^−5940/T	Assumed same as k(CH₃O+CO)
21	CH₃S+H₂O₂→CH₃SH+H₂O	3.01·10⁻¹³	Turnipseed et al.,1996
22	CH₃S+HCS→CH₃SH+CS	1.18·10⁻¹²×e^−910/T×(T/300)^0.65	Liu et al.,2006
23	CH₃S+HS→CH₃SH+S	1.66·10⁻¹¹	Assumed same as k(CH₃+HS)
24	CH₃SH+CH₃→CH₄+CH₃S	2.99·10⁻³¹	Kerr and Trotman-Dickenson, 1957
25	C₂H₆S+O→CH₃+CH₃+SO	1.30·10⁻¹¹×e^−410/T×(T/298)^1.1	Sander et al.,2006
26	CH₃SH+O→CH₃+HSO	1.30·10⁻¹¹×e^−410/T×(T/298)^1.1	Assumed same as k(C₂H₆S+O)
27	C₂H₆S₂+O→CH₃+CH₃S+SO	3.90·10⁻¹¹×e^290/T×(T/298)^1.1	Sander et al.,2006
28		1.10·10⁻¹¹×e^−240/T×(T/298)^1.1	Sander et al.,2006
29	C₂H₆S₂+OH→CH₃+CH₃SH+SO	6.00·10⁻¹¹×e^400/T×(T/298)^1.2	Sander et al.,2006
30	CH₃SH+OH→CH₃S+H₂O	9.90·10⁻¹²×e^360/T×(T/298)^1.07	Sander et al.,2006
31	C₂H₆S+O→CH₃+CH₃+SO	1.30·10⁻¹¹×e^−410/T×(T/298)^1.1	Sander et al.,2006
32	CH₃SH+O→CH₃+HSO	1.30·10⁻¹¹×e^−410/T×(T/298)^1.1	Assumed same as k(C₂H₆S+O)
33	CH₃SH+H→CH₃+H₂S	1.5·10⁻¹¹×e^−840/T	Amano et al.,1983
34	CH₃SH+H→H₂+CH₃S	4.82·10⁻¹¹×e^−1310/T	Amano et al.,1983
35	CH₃SH+OH→H₂O+CH₃S	9.9·10⁻¹²×e^360/T	DeMore and Yung, 1982
36	CH+CS₂→HCS+CS	3.49·10⁻¹⁰×e^−40/T	Zabarnick et al.,1989
37	CS+HS→CS₂+H	1.5·10⁻¹³×(1+0.6×den)	Assumed same as k(CO+OH)
38	CS+O→CO+S	2.7·10⁻¹⁰×e^−760/T	Atkinson et al.,2004
39	CS+O₂→CO+SO	5·10⁻²⁰	Wine et al.,1981
40	CS+O₂→OCS+O	4·10⁻¹⁹	Wine et al.,1981
41	CS+O₃→CO+SO₂	3·10⁻¹²	Wine et al.,1981
42	CS+O₃→OCS+O₂	3·10⁻¹²	Wine et al.,1981
43	CS+O₃→SO+CO₂	3·10⁻¹²	Wine et al.,1981
44	CS₂+O→CO+S₂	5.81·10⁻¹⁴	Singleton and Cvetanovic, 1988
45	CS₂+O→OCS+S	3·10⁻¹²×e^−650/T	Toon et al.,1987
46	CS₂+O→SO+CS	3.2·10⁻¹¹×e^−650/T	Toon et al.,1987
47	CS₂+OH→OCS+HS	2·10⁻¹⁵	Atkinson et al.,2004
48	CS₂+S→CS+S₂	1.9·10⁻¹⁴×e^−580/T×(T/300)^3.97	Woiki and Roth, 1995
49	CS₂+SO→OCS+S₂	2.4·10⁻¹³×e^−2370/T	Assumed same as k(SO^*+O₂)
50		1·10⁻¹²	Assumed same as
51		2.5·10⁻¹¹	Wine et al.,1981
52		1·10⁻¹²	Wine et al.,1981
53	C+HS→CS+H	4·10⁻¹¹	Assumed same as k(C+OH)
54	C+S₂→CS+S	3.3·10⁻¹¹	Assumed same as k(C+O₂)
55	C₂+S→C+CS	5·10⁻¹¹	Assumed same as k(C₂+O)
56	C₂+S₂→CS+CS	1.5·10⁻¹¹×e^−550/T	Assumed same as k(C₂+O₂)
57	CH+S→CS+H	9.5·10⁻¹¹	Assumed same as k(CH+CS₂)
58	CH+S₂→CS+HS	5.9·10⁻¹¹	Assumed same as k(CH+O₂)
59		3·10⁻¹¹	Assumed same as
60	CH₃+HCS→CH₄+CS	8.2·10⁻¹¹	Assumed same as k(CH₃+HCO)
61	H+CS+M→HCS+M	2.0·10⁻³³×e^−850/T×den	Assumed same as k(H+CO)
62	H+HCS→H₂+CS	1.2·10⁻¹⁰	Assumed same as k(H+HCO)
63	HS+CO→OCS+H	4.2·10⁻¹⁴×e^−7650/T	Kurbanov and Mamedov, 1995
64	HS+HCS→H₂S+CS	5.0·10⁻¹¹	Assumed same as k(HS+HCO)
65	OCS+CH→CO+HCS	1.99·10⁻¹⁰×e^−190/T	Zabarnick et al.,1989
66	S+CO+M→OCS+M	6.5·10⁻³³×e^−2180/T×den	Assumed same as k(CO+O)
67	S+HCS→H+CS₂	1.0·10⁻¹⁰	Assumed same as k(O+HCO→H+CO₂)
68	S+HCS→HS+CS	5.0·10⁻¹¹	Assumed same as k(O+HCO→HS+CO)
69		5.3·10⁻¹¹	Braun et al.,1970
70		k₀=8.75·10⁻³¹×e^524/T	Zahnle, 1986
		k_∞=8.3·10⁻¹
71	C+O₂→CO+O	3.3·10⁻¹¹	Donovan and Husain, 1970
72	C+OH→CO+H	4·10⁻¹¹	Giguere and Huebner, 1978
73	C₂+CH₄→C₂H+CH₃	5.05·10⁻¹¹×e^−297/T	Pitts et al.,1982
74	C₂+H₂→C₂H+H	1.77·10⁻¹⁰×e^−1469/T	Pitts et al.,1982
75	C₂+O→C+CO	5·10⁻¹¹	Prasad and Huntress, 1980
76	C₂+O₂→CO+CO	1.5·10⁻¹¹×e^−550/T	Baughcum and Oldenborg, 1984
77	C₂H+C₂H₂→HCAER+H	1.5·10⁻¹⁰	Stephens et al.,1987
78	C₂H+C₂H₆→C₂H₂+C₂H₅	3.6·10⁻¹¹	Lander et al.,1990
79	C₂H+C₃H₈→C₂H₂+C₃H₇	1.4·10⁻¹¹	Okabe, 1983
80	C₂H+CH₂CCH₂→HCAER+H	1.5·10⁻¹⁰	Pavlov et al.,2001
81	C₂H+CH₄→C₂H₂+CH₃	6.94·10⁻¹²×e^−250/T	Allen et al.,1992; Lander et al.,1990
82	C₂H+H+M→C₂H₂+M	k₀=2.64·10⁻²⁶×e^−721/T×(T/300)^−3.1	Tsang and Hampson, 1986
		k_∞=3.0·10⁻¹⁰
83	C₂H+H₂→C₂H₂+H	5.58·10⁻¹¹×e^−1443/T	Allen et al.,1992; Stephens et al.,1987
84	C₂H+O→CO+CH	1·10⁻¹⁰×e^−250/T	Zahnle, 1986
85	C₂H+O₂→CO+HCO	2·10⁻¹¹	Brown and Laufer, 1981
86	C₂H₂+H+M→C₂H₃+M	k₀=2.6·10⁻³¹	Romani et al.,1993
		k_∞=8.3·10⁻¹¹×e^−1374/T
87		2.9·10⁻¹¹×e^−1600/T	Zahnle, 1986
88	C₂H₂+OH+M→C₂H₂OH+M	k₀=5.5·10⁻³⁰	Sander et al.,2006
		k_∞=8.3·10⁻¹³×(T/300)⁻²
89	C₂H₂+OH+M→CH₂CO+H+M	k₀=5.8·10⁻³¹×e^1258/T	Perry and Williamson, 1982
		k_∞=1.4·10⁻¹²×e^388/T
90	C₂H₂+OH→CO+CH₃	2·10⁻¹²×e^−250/T	Hampson and Garvin, 1977
91	C₂H₂OH+H→H₂+CH₂CO	3.3·10⁻¹¹×e^−2000/T	Miller et al.,1982
92	C₂H₂OH+H→H₂O+C₂H₂	5·10⁻¹¹	Miller et al.,1982
93	C₂H₂OH+O→OH+CH₂CO	3.3·10⁻¹¹×e^−2000/T	Miller et al.,1982
94	C₂H₂OH+OH→H₂O+CH₂CO	1.7·10⁻¹¹×e^−1000/T	Miller et al.,1982
95	C₂H₃+C₂H₃→C₂H₄+C₂H₂	2.4·10⁻¹¹	Fahr et al.,1991
96	C₂H₃+C₂H₅→C₂H₄+C₂H₄	3·10⁻¹²	Laufer et al.,1983
97	C₂H₃+C₂H₅+M→CH₃+C₃H₅+M	k₀=1.9·10⁻²⁷	Romani et al.,1993
		k_∞=2.5·10⁻¹¹
98	C₂H₃+C₂H₆→C₂H₄+C₂H₅	3·10⁻¹³×e^−5170/T	Kasting et al.,1983
99	C₂H₃+CH₃→C₂H₂+CH₄	34·10⁻¹¹	Fahr et al.,1991
100	C₂H₃+CH₃+M→C₃H₆+M	k₀=1.3·10⁻²²	Raymond et al.,2006
		k_∞=1.2·10⁻¹⁰
101	C₂H₃+CH₄→C₂H₄+CH₃	2.4·10⁻²⁴×e^−2754/T×T^4.02	Tsang and Hampson, 1986
102	C₂H₃+H→C₂H₂+H₂	3.3·10⁻¹¹	Warnatz, 1984
103	C₂H₃+H₂→C₂H₄+H	2.6·10⁻¹³×e^−2646/T	Allen et al.,1992
104	C₂H₃+O→CH₂CO+H	5.5·10⁻¹¹	Hoyermann et al.,1981
105	C₂H₃+OH→C₂H₂+H₂O	8.3·10⁻¹²	Benson and Haugen, 1967
106	C₂H₄+H+M→C₂H₅+M	k₀=2.15·10⁻²⁹×e^−349/T	Lightfoot and Pilling, 1987
		k_∞=4.95·10⁻¹¹×e^−1051/T
107	C₂H₄+O→HCO+CH₃	5.5·10⁻¹²×e^−565/T	Hampson and Garvin, 1977
108	C₂H₄+OH+M→C₂H₄OH+M	k₀=1.0·10⁻²⁸×(T/300)^4.5	Sander et al.,2006
		k_∞=8.8·10⁻¹²×(T/300)^0.85
109	C₂H₄+OH→H₂CO+CH₃	2.2·10⁻¹²×e^385/T	Hampson and Garvin, 1977
110	C₂H₄OH+H→H₂+CH₃CHO	3.3·10⁻¹¹×e^−2000/T	Zahnle and Kasting, 1986
111	C₂H₄OH+H→H₂O+C₂H₄	5·10⁻¹¹	Miller et al.,1982
112	C₂H₄OH+O→OH+CH₃CHO	3.3·10⁻¹¹×e^−2000/T	Zahnle and Kasting, 1986
113	C₂H₄OH+OH→H₂O+CH₃CHO	1.7·10⁻¹¹×e^−1000/T	Zahnle and Kasting, 1986
114	C₂H₅+C₂H₃→C₂H₆+C₂H₂	6·10⁻¹²	Laufer et al.,1983
115	C₂H₅+C₂H₅→C₂H₆+C₂H₄	2.3·10⁻¹²	Tsang and Hampson, 1986
116	C₂H₅+CH₃→C₂H₄+CH₄	1.88·10⁻¹²×(T/300)^−0.5	Romani et al.,1993
117	C₂H₅+CH₃+M→C₃H₈+M	k₀=3.9·10⁻¹⁰×(T/300)^2.5	Romani et al.,1993
		k_∞=1.4·10⁻⁸×(T/300)^0.5
118	C₂H₅+H→C₂H₄+H₂	3·10⁻¹²	Tsang and Hampson, 1986
119	C₂H₅+H+M→C₂H₆+M	k₀=5.5·10⁻²³×e^−1040/T	Gladstone et al.,1996
		k_∞=1.5·10⁻¹⁰
120	C₂H₅+H→CH₃+CH₃	7.95·10⁻¹¹	Gladstone et al.,1996
121	C₂H₅+HCO→C₂H₆+CO	5·10⁻¹¹	Pavlov et al.,2001
122	C₂H₅+HNO→C₂H₆+NO	3·10⁻¹⁴	Pavlov et al.,2001
123	C₂H₅+O→CH₃+HCO+H	1.1·10⁻¹⁰	Pavlov et al.,2001
124	C₂H₅+O→CH₃CHO+H	1.33·10⁻¹⁰	Tsang and Hampson, 1986
125	C₂H₅+O₂+M→CH₃+HCO+OH+M	k₀=1.5·10⁻²⁸×(T/300)^3.0	Sander et al.,2006
		k_∞=8·10⁻¹²
126	C₂H₅+OH→CH₃CHO+H₂	1·10⁻¹⁰	Pavlov et al.,2001
127	C₂H₅+OH→C₂H₄+H₂O	4.0·10⁻¹¹	Pavlov et al.,2001
128	C₂H₆+O→C₂H₅+OH	8.62·10⁻¹²×e^−2920/T×(T/300)^1.5	Baulch et al.,1994
129	C₂H₆+O¹D→C₂H₅+OH	6.29·10⁻¹⁰	Matsumi et al.,1993
130	C₂H₆+OH→C₂H₅+H₂O	8.54·10⁻¹²×e^−1070/T	Sander et al.,2006
131	C₃H₂+H+M→C₃H₃+M	k₀=1.7·10⁻²⁶	Yung et al.,1984
		k_∞=1.5·10⁻¹⁰
132	C₃H₃+H+M→CH₂CCH₂+M	k₀=1.7·10⁻²⁶	Yung et al.,1984
		k_∞=1.5·10⁻¹⁰
133	C₃H₃+H+M→CH₃C₂H+M	k₀=1.7·10⁻²⁶	Yung et al.,1984
		k_∞=1.5·10⁻¹⁰
134	C₃H₅+CH₃→CH₂CCH₂+CH₄	4.5·10⁻¹²	Yung et al.,1984
135	C₃H₅+CH₃→CH₃C₂H+CH₄	4.5·10⁻¹²	Yung et al.,1984
136	C₃H₅+H+M→C₃H₆+M	k₀=1.0·10⁻²⁸	Yung et al.,1984
		k_∞=1.0·10⁻¹¹
137	C₃H₅+H→CH₂CCH₂+H₂	1.5·10⁻¹¹	Yung et al.,1984
138	C₃H₅+H→CH₃C₂H+H₂	1.5·10⁻¹¹	Yung et al.,1984
139	C₃H₅+H→CH₄+C₂H₂	1.5·10⁻¹¹	Yung et al.,1984
140	C₃H₆+H+M→C₃H₇+M	k₀=2.15·10⁻²⁹×e^−349/T	Pavlov et al.,2001
		k_∞=4.95·10⁻¹¹×e^−1051/T	Assumed same as k(C₂H₄+H)
141	C₃H₆+O→CH₃+CH₃CO	4.1·10⁻¹²×e^−38/T	Hampson and Garvin, 1977
142	C₃H₆+OH→CH₃CHO+CH₃	4.1·10⁻¹²×e^540/T	Hampson and Garvin, 1977
143	C₃H₇+CH₃→C₃H₆+CH₄	2.5·10⁻¹²×e^−200/T	Yung et al.,1984
144	C₃H₇+H→CH₃+C₂H₅	7.95·10⁻¹¹×e^−127/T	Pavlov et al.,2001
145	C₃H₇+O→C₂H₅CHO+H	1.1·10⁻¹⁰	Pavlov et al.,2001
146	C₃H₇+OH→C₂H₅CHO+H₂	1.1·10⁻¹⁰	Pavlov et al.,2001
147	C₃H₈+O+M→C₃H₇+OH+M	k₀=1.6·10⁻¹¹×e^−2900/T	Hampson and Garvin, 1977
		k_∞=2.2·10⁻¹¹×e^−2200/T
148	C₃H₈+O¹D→C₃H₇+OH	1.4·10⁻¹⁰	Pavlov et al.,2001
149	C₃H₈+OH→C₃H₇+H₂O	8.6·10⁻¹²×e^−615/T	Sander et al.,2006
150	CH+C₂H₂+M→C₃H₂+H+M	k₀=2.15·10⁻²⁹×e^−349/T	Romani et al.,1993
		k_∞=4.95·10⁻¹¹×e^−1051/T
151	CH+C₂H₄+M→CH₂CCH₂+H+M	k₀=1.75·10⁻¹⁰×e^61/T	Romani et al.,1993
		k_∞=5.3·10⁻¹⁰
152	CH+C₂H₄+M→CH₃C₂H+H+M	k₀=1.75·10⁻¹⁰×e^61/T	Romani et al.,1993
		k_∞=5.3·10⁻¹⁰
153	CH+CH₄+M→C₂H₄+H+M	k₀=2.5·10⁻¹¹×e^200/T	Romani et al.,1993
		k_∞=1.7·10⁻¹⁰
154	CH+CO₂→HCO+CO	5.9·10⁻¹²×e^−350/T	Berman et al.,1982
155	CH+H→C+H₂	1.4·10⁻¹¹	Becker et al.,1989
156		2.38·10⁻¹⁰×e^−1760/T	Zabarnick et al.,1986
157	CH+H₂+M→CH₃+M	k₀=8.75·10⁻³¹×e^524/T	Romani et al.,1993
		k_∞=8.3·10⁻¹¹
158	CH+O→CO+H	9.5·10⁻¹¹	Messing et al.,1981
159	CH+O₂→CO+OH	5.9·10⁻¹¹	Butler et al.,1981
160		7.14·10⁻¹²×e^−5050/T	Böhland et al.,1985
161		1·10⁻¹²	Zahnle, 1986
162		1.26·10⁻¹¹	Romani et al.,1993
163		5·10⁻¹⁵	Tsang and Hampson, 1986
164		8.8·10⁻¹²	Ashfold et al.,1981
165		3·10⁻¹¹	Ashfold et al.,1981
166		k₀=3.8·10⁻²⁵	Laufer, 1981; Laufer et al.,1983
		k_∞=3.7·10⁻¹²
167		k₀=3.8·10⁻²⁵	Laufer, 1981; Laufer et al.,1983
		k_∞=2.2·10⁻¹²
168		3·10⁻¹¹	Tsang and Hampson, 1986
169		3·10⁻¹¹	Tsang and Hampson, 1986
170		7·10⁻¹¹	Tsang and Hampson, 1986
171		k₀=1.0·10⁻²⁸	Yung et al.,1984
		k_∞=1.0·10⁻¹⁵
172		3.9·10⁻¹⁴	Laufer, 1981
173		4.7·10⁻¹⁰×e^−370/T	Zabarnick et al.,1986
174		k₀=3.1·10⁻³⁰×e^457/T	Gladstone et al.,1996
		k_∞=1.5·10⁻¹⁰
175		8·10⁻¹²	Huebner and Giguere, 1980
176		8.3·10⁻¹¹	Homann and Wellmann, 1983
177		1·10⁻¹¹	Huebner and Giguere, 1980
178		4.1·10⁻¹¹×e ^−750/T	Baulch et al.,1994
179	CH₂CCH₂+H→C₃H₅	k₀=8.9·10⁻²⁹×e^−1225/T×(T/300)^−2.0	Yung et al.,1984
		k_∞=1.4·10⁻¹¹×e^−1000/T
180	CH₂CCH₂+H→CH₃+C₂H₂	k₀=8.9·10⁻²⁹×e^−1225/T×(T/300)^−2.0	Yung et al.,1984
		k_∞=9.7·10⁻¹³×e^−1550/T
181	CH₂CCH₂+H→CH₃C₂H+H	1·10⁻¹¹×e^−1000/T	Yung et al.,1984
182	CH₂CO+H→CH₃+CO	1.9·10⁻¹¹×e^−1725/T	Michael et al.,1979
183	CH₂CO+O→H₂CO+CO	3.3·10⁻¹¹	Lee, 1980; Miller et al.,1982
184	CH₃+C₂H₃→C₃H₅+H	2.4·10⁻¹³	Romani et al.,1993
185	CH₃+CH₃+M→C₂H₆+M	k₀=4.0·10⁻²⁴×e^−1390/T×(T/300)^−7.0	Wagner and Wardlaw, 1988
		k_∞=1.79·10⁻¹⁰×e^−329/T
186	CH₃+CO+M→CH₃CO+M	1.4·10⁻³²×e^−3000/T×den	Watkins and Word, 1974
187	CH₃+H+M→CH₄+M	k₀=6.0·10⁻²⁸×(T/298)^−1.80	Baulch et al.,1994; Tsang and Hampson, 1986
		k_∞=2.0·10⁻¹⁰×(T/298)^−0.40
188	CH₃+H₂CO→CH₄+HCO	1.60·10⁻¹⁶×e^899/T×(T/298)^6.10	Baulch et al.,1994
189	CH₃+HCO→CH₄+CO	2.01·10⁻¹⁰	Tsang and Hampson, 1986
190	CH₃+HNO→CH₄+NO	1.85·10⁻¹¹×e^−176/T×(T/298)^0.6	Choi and Lin, 2005
191	CH₃+O→H₂CO+H	1.1·10⁻¹⁰	Sander et al.,2006
192	CH₃+O₂→H₂CO+OH	k₀=4.0·10⁻³¹×(T/300)^−3.6	Sander et al.,2006
		k_∞=1.2·10⁻¹²×(T/300)^−1.1
193	CH₃+O₃→H₂CO+HO₂	5.4·10⁻¹²×e^−220/T	Sander et al.,2006
194	CH₃+OH→CH₃O+H	9.3·10⁻¹¹×e^−1606/T×(T/298)	Jasper et al.,2007
195	CH₃+OH→CO+H₂+H₂	6.7·10⁻¹²	Fenimore, 1969
196	CH₃C₂H+H+M→C₃H₅+M	k₀=8.88·10⁻²⁹×e^−1225/T×(T/300)⁻²	Yung et al.,1984
		k_∞=9.7·10⁻¹²×e^−1550/T
197	CH₃C₂H+H→CH₃+C₂H₂	k₀=8.88·10⁻²⁹×e^−1225/T×(T/300)⁻²	Whytock et al.,1976
		k_∞=9.7·10⁻¹²×e^−1550/T
198	CH₃CHO+CH₃→CH₃CO+CH₄	2.8·10⁻¹¹×e^−1540/T	Zahnle, 1986
199	CH₃CHO+H→CH₃CO+H₂	2.8·10⁻¹¹×e^−1540/T	Zahnle, 1986
200	CH₃CHO+O→CH₃CO+OH	5.8·10⁻¹³	Washida, 1981
201	CH₃CHO+OH→CH₃CO+H₂O	1.6·10⁻¹¹	Niki et al.,1978
202	CH₃CO+CH₃→C₂H₆+CO	5.4·10⁻¹¹	Adachi et al.,1981
203	CH₃CO+CH₃→CH₄+CH₂CO	8.6·10⁻¹¹	Adachi et al.,1981
204	CH₃CO+H→CH₄+CO	1·10⁻¹⁰	Zahnle, 1986
205	CH₃CO+O→H₂CO+HCO	5·10⁻¹¹	Zahnle, 1986
206	CH₃O+CO→CH₃+CO₂	2.6·10⁻¹¹×e^−5940/T	Wen et al.,1989
207	CH₃O₂+H→CH₄+O₂	1.6·10⁻¹⁰	Tsang and Hampson, 1986
208	CH₃O₂+H→H₂O+H₂CO	1·10⁻¹¹	Zahnle et al.,2006
209	CH₃O₂+O→H₂CO+HO₂	1·10⁻¹¹	Vaghjiani and Ravishankara, 1990
210	CH₄+HS→CH₃+H₂S	2.99·10⁻³¹	Kerr and Trotman-Dickenson, 1957
211	CH₄+O→CH₃+OH	8.75·10⁻¹²×e^−4330/T×(T/298)^1.5	Tsang and Hampson, 1986
212	CH₄+O¹D→CH₃+OH	1.28·10⁻¹⁰	Sander et al.,2006
213	CH₄+O¹D→H₂CO+H₂	2.25·10⁻¹¹	Sander et al.,2006
214	CH₄+OH→CH₃+H₂O	2.45·10⁻¹²×e^−1775/T	Sander et al.,2006
215	CO+O+M→CO₂+M	1.7·10⁻³³×e^−1515/T×den	Tsang and Hampson, 1986
216	CO+OH→CO₂+H	1.5·10⁻¹³×(1+0.6×den)	Sander et al.,2006
217	H+CO+M→HCO+M	5.29·10⁻³⁴×e^−100/T×den	Baulch et al.,1994
218	H+H+M→H₂+M	8.85·10⁻³³×(T/298)^−0.6×den	Baulch et al.,1994
219	H+HCO→H₂+CO	1.5·10⁻¹⁰	Baulch et al.,1992
220	H+HNO→H₂+NO	3.01·10⁻¹¹×e^500/T	Tsang and Herron, 1991
221	H+HO₂→H₂+O₂	6.9·10⁻¹²	Sander et al.,2006
222	H+HO₂→H₂O+O	1.62·10⁻¹²	Sander et al.,2006
223	H+HO₂→OH+OH	7.29·10⁻¹¹	Sander et al.,2006
224	H+NO+M→HNO+M	2.1·10⁻³²×(T/298)^1.00×den	Hampson and Garvin, 1977
225	H+O₂+M→HO₂+M	5.7·10⁻³²×7.5·10⁻¹¹×(T/298)^1.6	Sander et al.,2006
226	H+O₃→OH+O₂	1.4·10⁻¹⁰×e^−470/T	Sander et al.,2006
227	H+OH+M→H₂O+M	6.8·10⁻³¹×(T/300)⁻²×den	McEwan and Phillips, 1975
228	H+SO+M→HSO+M	k₀=5.7·10⁻³²×(T/298)^1.6	Kasting, 1990
		k_∞=7.5·10⁻¹¹
229	H₂+O→OH+H	1.34·10⁻¹⁵×e^−1460/T×(T/298)^6.52	Robie et al.,1990
230	H₂+O¹D→OH+H	1.1·10⁻¹¹	Sander et al.,2006
231	H₂+OH→H₂O+H	5.5·10⁻¹²×e^−2000/T	Sander et al.,2006
232	H₂CO+H→H₂+HCO	2.14⁻¹²×e^−1090/T×(T/298)^1.62	Baulch et al.,1994
233	H₂CO+O→HCO+OH	3.4·10⁻¹¹×e^−1600/T	Sander et al.,2006
234	H₂CO+OH→H₂O+HCO	5.5·10⁻¹²×e^125/T	Sander et al.,2006
235	H₂O+O¹D→OH+OH	2.2·10⁻¹⁰	Sander et al.,2006
236	H₂O₂+O→OH+HO₂	1.4·10⁻¹²×e^−2000/T	Sander et al.,2006
237	H₂O₂+OH→HO₂+H₂O	2.9·10⁻¹²×e^−160/T	Sander et al.,2006
238	H₂S+H→H₂+HS	3.66·10⁻¹²×e^−455/T×(T/298)^1.94	Peng et al.,1999
239	H₂S+O→OH+HS	9.2·10⁻¹²×e^−1800/T	Sander et al.,2006
240	H₂S+OH→H₂O+HS	6.0·10⁻¹²×e^−70/T	Sander et al.,2006
241	HCO+H+M→CO+M	6.0·10⁻¹¹×e^−7721/T×den	Krasnoperov et al.,2004
242	HCO+H₂CO→CH₃O+CO	3.8·10⁻¹⁷	Wen et al.,1989
243	HCO+HCO→H₂CO+CO	3.0·10⁻¹¹	Tsang and Hampson, 1986
244	HCO+NO→HNO+CO	1.2·10⁻¹¹	Tsang and Hampson, 1986
245	HCO+O₂→HO₂+CO	5.2·10⁻¹²	Sander et al.,2006
246	HNO+NO+M→H+M	1.04·10⁻⁶×e^25618/T×(T/298)^−1.61×den	Tsang and Hampson, 1986
247	HNO₂+OH→H₂O+NO₂	1.8·10⁻¹¹×e^−390/T	Sander et al.,2006
248	HNO₃+OH→H₂O+NO₂+O	7.2·10⁻¹⁵×e^−785/T+	Sander et al.,2006
		(1.9·10⁻³³×e^725/T×den)/
		(1+4.6·10⁻¹⁶×e^−715/T×den)
249	HO₂+HO₂→H₂O₂+O₂	k₀=2.3·10⁻¹³×e^590/T	Sander et al.,2006
		k_∞=1.7·10⁻³³×e^1000/T
250	HO₂+O→OH+O₂	3.0·10⁻¹¹×e^200/T	Sander et al.,2006
251	HO₂+O₃→OH+O₂+O₂	1.1·10⁻¹⁴×e^−490/T	Sander et al.,2006
252	HS+H→H₂+S	3.0·10⁻¹¹	Schofield, 1973
253	HS+H₂CO→H₂S+HCO	1.7·10⁻¹¹×e^−800/T	Sander et al.,2006
254	HS+HCO→H₂S+CO	5.0·10⁻¹¹	Kasting, 1990
255	HS+HO₂→H₂S+O₂	1.0·10⁻¹¹	Stachnik and Molina, 1987
256	HS+HS→H₂S+S	1.5⁻¹¹	Schofield, 1973
257	HS+NO₂→HSO+NO	2.9·10⁻¹¹×e^240/T	Sander et al.,2006
258	HS+O→H+SO	1.6·10⁻¹⁰	Sander et al.,2006
259	HS+O₂→OH+SO	4.0·10⁻¹⁹	Sander et al.,2006
260	HS+O₃→HSO+O₂	9.0·10⁻¹²×e^−280/T	Sander et al.,2006
261	HS+S→H+S₂	2.2·10⁻¹¹×e^−120/T	Kasting, 1990
262	HSO+H→H₂+SO	6.48·10⁻¹²	Sander et al.,2006
263	HSO+H→HS+OH	7.29·10⁻¹¹	Sander et al.,2006
264	HSO+HS→H₂S+SO	1·10⁻¹²	Kasting, 1990
265	HSO+NO→HNO+SO	1.0·10⁻¹⁵	Atkinson et al.,2004
266	HSO+O→OH+SO	3.0·10⁻¹¹×e^−200/T	Kasting, 1990
267	HSO+OH→H₂O+SO	5.2·10⁻¹²	Sander et al.,2006
268	HSO+S→HS+SO	1·10⁻¹¹	Kasting, 1990
269	HSO₃+H→H₂+SO₃	1.0·10⁻¹¹	Kasting, 1990
270	HSO₃+O→OH+SO₃	1.0·10⁻¹¹	Kasting, 1990
271	HSO₃+O₂→HO₂+SO₃	1.3·10⁻¹²×e^−330/T	Sander et al.,2006
272	HSO₃+OH→H₂O+SO₃	1.0·10⁻¹¹	Kasting, 1990
273	N+NO→N₂+O	2.1·10⁻¹¹×e^−100/T	Sander et al.,2006
274	N+O₂→NO+O	1.5·10⁻¹²×e^−3600/T	Sander et al.,2006
275	N+OH→NO+H	3.8·10⁻¹¹×e^85/T	Atkinson et al.,1989
276	N₂H₃+H→NH₂+NH₂	2.7·10⁻¹²	Gehring et al.,1971
277	N₂H₃+N₂H₃→N₂H₄+N₂+H₂	6·10⁻¹¹	Kuhn and Atreya, 1979
278	N₂H₄+H→N₂H₃+H₂	9.9·10⁻¹²×e^−1200/T	Stief and Payne, 1976
279	NH+H+M→NH₂+M	(6·10⁻³⁰×den)/(1+3·10⁻²⁰×den)	Kasting, 1982
280	NH+NO→N₂+OH	4.9·10⁻¹¹	Sander et al.,2006
281	NH+O→N+OH	1·10⁻¹¹	Kasting, 1982
282	NH+O→NH₂+CO	1·10⁻¹¹	Pavlov et al.,2001
283	NH₂+H+M→NH₃+M	(6·10⁻³⁰×den)/(1+3·10⁻²⁰×den)	Gordon et al.,1971
284	NH₂+HCO→NH₃+CO	1·10⁻¹¹	Pavlov et al.,2001
285	NH₂+NH₂→N₂H₄	1·10⁻¹⁰	Gordon et al.,1971
286	NH₂+NO→N₂+H₂O	3.8·10⁻¹²×e^450/T	Sander et al.,2006
287	NH₂+O→HNO+H	5·10⁻¹²	Albers et al.,1969
288	NH₂+O→NH+OH	5·10⁻¹²	Albers et al.,1969
289		3·10⁻¹¹	Kasting, 1982
290		3·10⁻¹¹	Kasting, 1982
291	NH₃+O¹D→NH₂+OH	2.5·10⁻¹⁰	Sander et al.,2006
292	NH₃+OH→NH₂+H₂O	1.7·10⁻¹²×e^−710/T	Sander et al.,2006
293	NO+HO₂→NO₂+OH	3.5·10⁻¹²×e^250/T	Sander et al.,2006
294	NO+O+M→NO₂+M	9·10⁻³¹×3·10⁻¹¹×(T/298)^1.5	Sander et al.,2006
295	NO+O₃→NO₂+O₂	2.0·10⁻¹²×e^−1400/T	Sander et al.,2006
296	NO+OH+M→HNO₂+M	k₀=7·10⁻³¹×(T/298)^2.6	Sander et al.,2006
		k_∞=3.6·10⁻¹¹×(T/298)^0.1
297	NO₂+H→NO+OH	4·10⁻¹⁰×e^−340/T	Sander et al.,2006
298	NO₂+O→NO+O₂	5.6·10⁻¹²×e^180/T	Sander et al.,2006
299	NO₂+OH+M→HNO₃+M	k₀=2.0·10⁻³⁰×(T/298)^3.0	Sander et al.,2006
		k_∞=2.5·10⁻¹¹
300	O+HCO→H+CO₂	5.0·10⁻¹¹	Tsang and Hampson, 1986
301	O+HCO→OH+CO	1.0·10⁻¹⁰	Hampson and Garvin, 1977
302	O+HNO→OH+NO	5.99·10⁻¹¹	Tsang and Hampson, 1986
303	O+O+M→O₂+M	9.46·10⁻³⁴×e^480/T×den	Campbell and Gray, 1973
304	O+O₂+M→O₃+M	6·10⁻³⁴×3·10⁻¹¹×(T/298)^2.40	Sander et al.,2006
305	O+O₃→O₂+O₂	8.0·10⁻¹²×e^−2060/T	Sander et al.,2006
306	O¹D+M→O+M	1.8·10⁻¹¹×e^110/T	Sander et al.,2006
307	O¹D+O₂→O+O₂	3.2·10⁻¹¹×e^70/T	Sander et al.,2006
308	OH+HCO→H₂O+CO	1.7·10⁻¹⁰	Baulch et al.,1992
309	OH+HNO→H₂O+NO	5·10⁻¹¹	Sun et al.,2001
310	OH+HO₂→H₂O+O₂	4.8·10⁻¹¹×e^250/T	Sander et al.,2006
311	OH+O→H+O₂	2.2·10⁻¹¹×e^120/T	Sander et al.,2006
312	OH+O₃→HO₂+O₂	1.6·10⁻¹²×e^−940/T	Sander et al.,2006
313	OH+OH→H₂O+O	4.2·10⁻¹²×e^−240/T	Sander et al.,2006
314	OH+OH→H₂O₂	6.9·10⁻³¹×1.5·10⁻¹¹×(T/298)^0.80	Sander et al.,2006
315	S+CO₂→SO+CO	1.0·10⁻²⁰	Yung and Demore, 1982
316	S+HCO→HS+CO	5.0·10⁻¹¹	Kasting, 1990
317	S+HO₂→HS+O₂	1.5·10⁻¹¹	Kasting, 1990
318	S+HO₂→SO+OH	1.5·10⁻¹¹	Kasting, 1990
319	S+O₂→SO+O	2.3·10⁻¹²	Sander et al.,2006
320	S+O₃→SO+O₂	1.2·10⁻¹¹	Sander et al.,2006
321	S+OH→SO+H	6.6·10⁻¹¹	Sander et al.,2006
322	S+S+M→S₂+M	1.98·10⁻³³×e^−206/T×den	Du et al.,2008
323	S+S₂+M→S₃+M	2.8·10⁻³²×den	Kasting, 1990
324	S+S₃+M→S₄+M	2.8·10⁻³¹×den	Kasting, 1990
325	S₂+O→S+SO	1.1·10⁻¹¹	Hills et al.,1987
326	S₂+S₂+M→S₄+M	2.8·10⁻³¹×den	Baulch et al.,1976
327	S₄+S₄+M→S₈AER+M	2.8·10⁻³¹×den	Kasting, 1990
328	SO+HCO→HSO+CO	5.6·10⁻¹²×(T/298)^−0.4	Kasting, 1990
329	SO+HO₂→SO₂+OH	2.8·10⁻¹¹	Kasting, 1990
330	SO+NO₂→SO₂+NO	1.4·10⁻¹¹	Sander et al.,2006
331	SO+O+M→SO₂+M	6.0·10⁻³¹×den	Sander et al.,2006
332	SO+O₂→O+SO₂	2.4·10⁻¹³×e^−2370/T	Sander et al.,2006
333	SO+O₃→SO₂+O₂	4.5·10⁻¹²×e^−1170/T	Atkinson et al.,2004
334	SO+OH→SO₂+H	8.6·10⁻¹¹	Sander et al.,2006
335	SO+SO→SO₂+S	3.5·10⁻¹⁵	Martinez and Herron, 1983
336	SO₂+HO₂→SO₃+OH	8.63·10⁻¹⁶	Lloyd, 1974
337	SO₂+O+M→SO₃+M	k₀=1.3·10⁻³³×(T/298)^−3.6	Sander et al.,2006
		k_∞=1.5·10⁻¹¹
338	SO₂+OH+M→HSO₃+M	k₀=3·10⁻³¹×(T/298)^3.3	Sander et al.,2006
		k_∞=1.5·10⁻¹²
339		1.0·10⁻¹¹	Turco et al.,1982
340		1.0·10⁻¹²	Turco et al.,1982
341		1.0·10⁻¹⁶	Turco et al.,1982
342		4.0·10⁻¹²	Turco et al.,1982
343		1.5·10⁻¹³	Turco et al.,1982
344		7.0·10⁻¹⁴	Turco et al.,1982
345	SO₃+H₂O→H₂SO₄	1.2·10⁻¹⁵	Sander et al.,2006
346	SO₃+SO→SO₂+SO₂	2.0·10⁻¹⁵	Chung et al.,1975
347		2.2·10⁺⁴	Turco et al.,1982
348		1.5·10⁺³	Turco et al.,1982
349		1.13·10⁺³	Turco et al.,1982
350	O₂+hν→O+O¹D	1.51·10⁺⁰²
351	O₂+hν→O+O	2.90·10⁺⁰⁰
352	H₂O+hν→H+OH	1.65·10⁻⁰¹
353	O₃+hν→O₂+O¹D	6.44·10⁻⁰⁴
354	O₃+hν→O₂+O	1.64·10⁻⁰⁴
355	H₂O₂+hν→OH+OH	2.79·10⁻¹⁴
356	CO₂+hν→CO+O	2.50·10⁺⁰¹
357	H₂CO+hν→H₂+CO	7.71·10⁻⁰¹
358	H₂CO+hν→HCO+H	9.33·10⁻⁰¹
359	CO₂+hν→CO+O¹D	2.73·10⁺⁰³
360	HO₂+hν→OH+O	0.00·10⁺⁰⁰
361		1.75·10⁺⁰⁰
362		0.00
363		1.48·10⁻⁰⁵
364	HNO₂+hν→NO+OH	8.68·10⁻²²
365	HNO₃+hν→NO₂+OH	2.74·10⁻²⁸
366	NO+hν→N+O	2.04·10⁻¹⁰
367	NO₂+hν→NO+O	4.40·10⁻¹⁴
368		6.67·10⁻⁰⁴
369	SO+hν→S+O	0.00·10⁺⁰⁰
370	SO₂+hν→SO+O	1.37·10⁻¹⁰
371	H₂S+hν→HS+H	1.00·10⁻²³
372		1.52·10⁻⁰⁹
373		8.14·10⁻¹³
374	S₂+hν→S+S	5.94·10⁻⁴²
375	S₂+hν→S₂	0.00·10⁺⁰⁰
376	H₂SO₄+hν→SO₂+OH+OH	1.66·10⁻¹³
377	SO₃+hν→SO₂+O	0.00·10⁺⁰⁰
378		9.70·10⁻¹¹
379		1.42·10⁻⁰⁹
380		9.78·10⁻¹¹
381	HSO+hν→HS+O	7.19·10⁻¹⁷
382	S₄+hν→S₂+S₂	0.00·10⁺⁰⁰
383	S₃+hν→S₂+S	4.22·10⁻⁷²
384	NH₃+hν→NH₂+H	6.00·10⁻³⁴
385	N₂H₄+hν→N₂H₃+H	9.75·10⁻⁹³
386	NH+hν→N+H	3.99·10⁻³⁵
387	NH₂+hν→NH+H	7.49·10⁻³⁷
388		3.99·10⁻³⁵
389		3.99·10⁻³⁵
390	C₂H₂+hν→C₂H+H	5.51·10⁻⁰⁷
391	C₂H₂+hν→C₂+H₂	4.09·10⁻⁰⁷
392	C₂H₄+hν→C₂H₂+H₂	5.51·10⁻⁰⁷
393	C₃H₈+hν→C₃H₆+H₂	1.45·10⁻¹²
394		2.49·10⁻¹³
395	C₃H₈+hν→C₂H₄+CH₄	1.08·10⁻¹²
396	C₃H₈+hν→C₂H₅+CH₃	5.88·10⁻¹³
397	C₂H₆+hν→C₂H₂+H₂+H₂	1.80·10⁻⁰⁵
398	C₂H₆+hν→C₂H₄+H+H	1.93·10⁻⁰⁵
399	C₂H₆+hν→C₂H₄+H₂	5.29·10⁻⁰⁷
400	C₂H₆+hν→CH₃+CH₃	4.79·10⁻⁰⁶
401	C₂H₄+hν→C₂H₂+H+H	5.29·10⁻⁰⁷
402	C₃H₆+hν→C₂H₂+CH₃+H	5.26·10⁻¹⁶
403		1.42·10⁺⁰⁰
404	CH₄+hν→CH₃+H	2.91·10⁺⁰⁰
405	CH+hν→C+H	9.52·10⁻⁰⁶
406		8.21·10⁻¹⁰
407	CH₃CHO+hν→CH₃+HCO	1.14·10⁻⁰⁸
408	CH₃CHO+hν→CH₄+CO	1.14·10⁻⁰⁸
409	C₂H₅CHO+hν→C₂H₅+HCO	6.42·10⁻⁰⁷
410	C₃H₃+hν→C₃H₂+H	6.88·10⁻⁰⁷
411	CH₃C₂H+hν→C₃H₃+H	6.42·10⁻⁰⁷
412	CH₃C₂H+hν→C₃H₂+H₂	2.41·10⁻⁰⁷
413	CH₃C₂H+hν→CH₃+C₂H	3.21·10⁻⁰⁸
414	CH₂CCH₂+hν→C₃H₃+H	6.49·10⁻¹³
415	CH₂CCH₂+hν→C₃H₂+H₂	2.43·10⁻¹³
416		9.73·10⁻¹⁴
417	C₃H₆+hν→CH₂CCH₂+H₂	8.81·10⁻¹⁶
418		3.09·10⁻¹⁷
419	C₃H₆+hν→C₂H+CH₄+H	1.43·10⁻¹⁰
420	OCS+hν→CO+S	2.67·10⁻³⁶
421	CS₂+hν→CS+S	5.40·10⁻⁴⁷
422	CH₃SH+hν→H+CH₃S	1.48·10⁻³⁰
423	CH₃SH+hν→HS+CH₃	1.11·10⁻³¹
424	C₂H₆S+hν→CH₃S+CH₃	4.01·10⁻⁹³
425	C₂H₆S₂+hν→CH₃S+CH₃S	1.65·10⁻³⁴
426		6.57·10⁻⁴⁸

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For photolysis reactions (bottom of table), the “Reaction rate constant” column shows the reaction rate (not the rate constant) at the top of the atmosphere during our “standard” simulation, the modern-day fluxes of CH₄, H₂S, and the S_org species on a planet orbiting the Sun. For more on how to calculate reaction rates, see Sander et al. (2006).

The model grid is composed of 100 plane-parallel layers that are each 1 km thick in altitude. We did not perform climate calculations for this work; instead, we assumed a temperature profile for an aerosol-free, ozone-free atmosphere. This profile had a surface temperature of 278 K that decreased to 180 K at the tropopause and was isothermal through the stratosphere. The relatively low surface temperature was picked for consistency with previous Archean photochemistry and climate models (Haqq-Misra et al., 2008), and the isothermal stratosphere is consistent with the model's lack of O₃. The code calculates the mixing ratios of each species in each layer by solving the coupled mass-continuity/flux equations with the reverse Euler method (appropriate for stiff systems) and a variable time-stepping algorithm. For further details on the photochemical code, see Pavlov et al. (2001) and references therein.

Unless otherwise stated, all model runs were for a 1-bar, N₂-dominated atmosphere with 3% CO₂ (30,000 ppmv, or ∼100 times the present level of CO₂ in Earth's atmosphere) and CH₄/CO₂ ratios < 0.1. These boundary conditions prevent formation of a significant organic haze (Pavlov et al., 2001; Trainer et al., 2006; Domagal-Goldman et al., 2008). These concentrations and the model's other chemical boundary conditions are by no means unique; however, they were chosen for consistency with a methanogen-acetogen ecosystem (Kharecha et al., 2005). The modeling of haze-free atmospheres is, from a photochemical standpoint, conservative. Including haze in the model would shield the gases we are studying from UV radiation and thereby increase their mixing ratios.

2.2. Boundary conditions

At the top of the atmosphere we allowed H and H₂ to escape at the diffusion-limited rate (Walker, 1977). We also applied a constant downward flux of CO and O at the top of our model atmosphere. This accounts for CO and O that is produced from CO₂ photolysis above the top layer of our atmosphere and subsequently flows downward into the model grid. For all other species, we used a zero-flux boundary condition at the top of the atmosphere (i.e., no escape).

At the bottom of the atmosphere, we used constant deposition velocities (to account for reactions with surface rocks and for dissolution in the ocean) for all species except the S_org species, CH₄, and NH₃. In addition to constant deposition velocities, H₂S, SO₂, and H₂ had volcanic fluxes of 1×10⁹ molecules/cm²/s, 1×10¹⁰ molecules/cm²/s, and 3×10¹⁰ molecules/cm²/s, respectively, consistent with past models of Archean Earth (Zahnle et al., 2006) that assume volcanism rates about 3 times modern-day values. These fluxes were distributed throughout the troposphere to simulate volcanism. CH₄ was modeled with a constant flux of 200 Tg C/year (7×10¹⁰ molecules/cm²/s) into the bottom layer of the atmosphere, in line with estimates of modern-day non-anthropogenic fluxes on Earth (Intergovernmental Panel on Climate Change, 2007). [The total CH₄ flux today is about 2 times higher; see the Intergovernmental Panel on Climate Change (2007)]. Despite this modern-day flux, the concentrations of CH₄ in our models were much higher than they are today because the lack of atmospheric O₂ allowed CH₄ to accumulate. We imposed a constant mixing ratio of 10⁻¹⁰ for NH₃. The corresponding surface flux needed to maintain this mixing ratio was 12.4 Tg N/year, slightly larger than the present-day non-anthropogenic NH₃ flux, 10.5 Tg N/year (Intergovernmental Panel on Climate Change, 2007). All photochemical boundary conditions are listed in Table 2.

Table 2.

A List of Species in Our Photochemical Code along with the Lower Boundary Condition Type and Values, the Latter Given in cgs Units: cm/s for Deposition Velocity (V_dep), Dimensionless Mixing Ratio by Volume for Fixed Concentration (f₀), and Molecules/cm²/s for Flux (flux)

Species	Lower boundary type	V_dep/f₀/flux
O	constant deposition velocity	1
O₂	constant deposition velocity	1·10⁻⁰⁴
H₂O	constant deposition velocity	0
H	constant deposition velocity	1
OH	constant deposition velocity	1
HO₂	constant deposition velocity	1
H₂O₂	constant deposition velocity	2·10⁻⁰¹
H₂	constant deposition velocity^*	2.4·10⁻⁰⁴
CO	constant deposition velocity	1.2·10⁻⁰⁴
HCO	constant deposition velocity	1
H₂CO	constant deposition velocity	2·10⁻⁰¹
CH₄	constant flux	7·10⁺¹⁰
CH₃	constant deposition velocity	1
C₂H₆	constant deposition velocity	0
NO	constant deposition velocity	3·10⁻⁰⁴
NO₂	constant deposition velocity	3·10⁻⁰³
HNO	constant deposition velocity	1
H₂S	constant deposition velocity^*	2·10⁻⁰²
HS	constant deposition velocity	1
S	constant deposition velocity	1
SO	constant deposition velocity	3·10⁻⁰⁴
SO₂	constant deposition velocity^*	1
H₂SO₄	constant deposition velocity	1
HSO	constant deposition velocity	1
S₂	constant deposition velocity	0
NH₃	constant mixing ratio	1·10⁻¹⁰
NH₂	constant deposition velocity	1
N₂H₃	constant deposition velocity	1
N₂H₄	constant deposition velocity	2·10⁻⁰¹
	constant deposition velocity	0
C₂H₅	constant deposition velocity	0
C₂H₂	constant deposition velocity	0
C₂H₄	constant deposition velocity	0
C₃H₈	constant deposition velocity	0
C₂H₃	constant deposition velocity	0
C₃H₆	constant deposition velocity	0
C₃H₂	constant deposition velocity	0
CH₂CCH₂	constant deposition velocity	0
CH₃C₂H	constant deposition velocity	0
C₂H₆S₂ (DMDS)	constant flux	0
C₂H₆S (DMS)	constant flux	4.20·10⁺⁰⁹
CH₃S	constant deposition velocity	1·10⁻⁰²
CH₃SH	constant flux	8.3·10⁺⁰⁸
CS₂	constant flux	1.4·10⁺⁰⁷
OCS	constant flux	1.4·10⁺⁰⁷
CS	constant deposition velocity	1·10⁻⁰⁴

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^{^*}

In addition to a constant deposition velocity, we also use a volcanic flux for these gases. Specifically, we used volcanic fluxes of 3·10¹⁰ molecules/cm²/s of H₂, 1·10¹⁰ molecules/cm²/s of SO₂, and 1·10⁹ molecules/cm²/s of H₂S.

We parameterized the biological production of S_org. The modern-day S_org fluxes, predominantly biological in source, are as follows (in units of molecules/cm²/s): 0 for DMDS, 4.2×10⁹ for DMS, 0 for CH₃S, 8.3×10⁸ for CH₃SH, 1.4×10⁷ for CS₂, 1.4×10⁷ for OCS, and 0 for CS (Kettle et al., 2001). We will use “MDF” as a unit to represent these modern-day fluxes in the rest of this paper, such that 1 MDF S_org is equivalent to an atmosphere that receives all S_org species at the above fluxes. DMDS, CH₃S, and CS have zero direct biological production but are produced photochemically from other S_org species and are needed to ensure a comprehensive modeling of S_org chemistry. To determine the effect of S_org fluxes on S_org mixing ratios and ultimately on disc-averaged planetary spectra, we parameterized S_org flux rates by holding the ratios of these fluxes constant and multiplying each flux by a common factor.

Most S_org species are produced via methylation of (addition of methyl groups to) CH₃SH or dehydrogenation of (removal of H atoms from) CH₃SH, or both. The main modern-day global source of CH₃SH is the degradation of methionine, an amino acid that contains a terminal methio group (−SCH₃), from eukaryotes. Based on the production rate of methionine, Pilcher (2003) estimated the flux of CH₃SH during the Archean to be ∼3×10⁹ mol/year, or about 0.01 MDF CH₃SH. This estimate agrees with what one would get by simply scaling CH₃SH production linearly with net primary productivity, as that is also estimated to have been ∼0.01 times the modern value (Kharecha et al., 2005). Because the Archean is our lone example of an anoxic planet, Pilcher's work serves as an estimate for the S_org fluxes on extrasolar planets with anoxic surface conditions. However, these fluxes could vary if methionine (or some other S-containing amino acid) was more or less prevalent in the planet's biota or if the biospheric productivity was different. Thus, in our primary suite of model runs, we parameterize the S_org fluxes from methionine degradation, using values from 0 to 3000 times those estimated by Pilcher (2003) (this is equivalent to 0–30 MDF S_org).

The direct production of CH₃SH for metabolic purposes could lead to higher S_org fluxes. Methanosarcina acetivorans, a methanogen, can produce CH₃SH via the metabolic reaction 3CO+H₂S+H₂O→CH₃SH+2CO₂ (Moran et al., 2008). In the rest of this manuscript, we will refer to this metabolism as “mercaptogenesis” and to the organisms that utilize it as “mercaptogens.” Assuming substrate-limited (CO-/H₂S-limited) conditions with no competition for substrates places an upper limit on mercaptogenesis. CO should build up to extremely high levels on planets with anoxic atmospheres unless consumed by biota (Zahnle, 1986; Kharecha et al., 2005); thus, H₂S is likely the limiting substrate on such planets. Estimates of the net primary productivity of S-consumers on Archean Earth vary by orders of magnitude, from 5×10⁹ mol S/year (Kharecha et al., 2005) to 2×10¹⁴ mol S/year (Canfield, 2005). Both estimates have caveats: the lower estimate did not include a complete S cycle that allowed for recycling of S, and the upper estimate neglected inorganic sinks for S such as metal-sulfide deposition. The former omission likely has a larger impact, so we used Canfield's estimate as an upper limit to S utilization. If mercaptogens accounted for all H₂S used by metabolism, the range of the above S consumption estimates would correspond to CH₃SH fluxes of ∼3×10⁹ to 1×10¹⁴ moles CH₃SH/year, or 0.03–1000 MDF CH₃SH. Thus, 1000 MDF CH₃SH is an upper limit to the CH₃SH produced by mercaptogens on early Earth. The actual CH₃SH production was likely much lower than this, due to competition for CO and H₂S from other metabolisms or from scavenging of S from the oceans by metal precipitates. On an extrasolar planet, the CH₃SH production rate could be higher if the planet has larger volcanic H₂S flux rates. Constraining such fluxes may be possible via absorption features of volcanic gases in planetary spectra (Kaltenegger et al., 2010).

Unfortunately, no anoxic ocean model currently exists that includes biological S recycling and a complete accounting of oceanic S sources. Furthermore, no code exists that can model mercaptogens in the context of CO-consuming methanogens and sulfur oxidizers that could compete for substrates. These problems might eventually be addressed by the development of ocean biogeochemistry codes with flexible chemistries and a wide variety of metabolisms. In the absence of such codes, we parameterized CH₃SH fluxes from 1 to 100 MDF to simulate a biosphere with CO-consuming mercaptogens. Because the CO they consume would otherwise be used by methanogens, we decreased the biological CH₄ flux in proportion to the biological CH₃SH flux in these simulations. In this set of mercaptogenesis experiments with 1–100 MDF CH₃SH, we held the fluxes of the other S_org gases (DMS, DMDS, OCS, and CS₂) constant at 1 MDF, because, unlike CH₃SH, these gases are not directly produced by this metabolism. To distinguish between the two sets of experiments, we label model simulations where we changed the flux of all S_org gases with “X MDF S_org,” and label model simulations where we changed only the flux of CH₃SH with “X MDF CH₃SH.”

Each set of S_org boundary conditions was applied to planets orbiting stars of three different spectral types, following Segura et al. (2005). Specifically, we used time-averaged spectra of the Sun, the active M dwarf AD Leo, and a model-generated M dwarf with a surface temperature of 3100 K and no chromosphere (Allard et al., 1997). This star, referred to as “T3100” in the remainder of this manuscript, is not presented as a physically meaningful case but rather as a low-UV-flux, end-member simulation. All stellar spectra were scaled such that the total energy flux at the top of the model planet's atmosphere was 1092 W/m², including radiation outside the bounds of our photochemical model wavelength grid. This is 80% of the flux Earth currently receives from the Sun, which is in line with the amount of energy the anoxic, Archean Earth received. Because the total energy flux received by the planet is the same, this is equivalent to assuming that the planet orbits within the habitable zone of that star. The resulting scaled stellar spectra are plotted in Fig. 1, as binned for use in the photochemical code.

FIG. 1. — (a) The stellar energy distribution at a planet receiving the same amount of total energy flux that the Earth received ∼2.5 billion years ago for three different stars: the Sun, AD Leo, and T3100 (a model M dwarf that has no chromosphere). (b) The bottom panel is an expansion of the UV region of the top panel, with a logarithmic y axis. The bottom panel also shows the absorption cross section of CH₃SH, units for which are on the right y axis (also logarithmic). Color images available online at www.liebertonline.com/ast

2.3. Radiative transfer code

We used the line-by-line Spectral Mapping Atmospheric Radiative Transfer model (Meadows and Crisp, 1996; Crisp, 1997) to generate synthetic planetary spectra of our model planets. Spectra were computed by using the vertical mixing ratio profiles of CH₃SH, DMS, DMDS, OCS, CS₂, SO₂, H₂S, CH₄, C₂H₆, CO₂, and H₂O generated by the photochemical code. The underlying surface consisted of a 278 K global ocean with an emissivity of ∼1, and we used the same assumed temperature structure applied in the photochemical model. The input stellar spectra and molecular absorption data were obtained from the Virtual Planetary Laboratory's online database (http://vpl.astro.washington.edu/spectra/VPLSpectra/frontpage.htm) and include molecular line parameters from the HITRAN (Rothman et al., 2005) and PNNL databases (Sharpe et al., 2004).

We did not include any aerosols in our spectral model, so the model spectra shown here should be considered idealized “clear sky” simulations. However, we limited parameter space (see Boundary conditions, above) so that the atmospheric CH₄/CO₂ ratio was less than 0.1, a condition for which thick organic haze layers will not form (Trainer et al., 2006; Haqq-Misra et al., 2008). S₈ and sulfate hazes were also limited by these conditions. Assuming Mie scattering, all S₈, hydrocarbon, and sulfate particles in our simulations had extinction optical depths less than 0.05 within the “IR window” between 8.5 and 13 μm in which most of the absorption features explored here appear. While organo-sulfate particles can form in sulfur-rich anoxic atmospheres (DeWitt et al., 2010), the optical properties of these particles have not yet been explored. Water clouds may also impact the spectra simulated here. For more on the effects of water clouds, see Robinson et al. (2011). We leave the exploration of aerosol and cloud effects for future studies.

3. Results

The habitable-zone planets around stars with lower surface temperatures receive proportionally fewer UV photons and more long-wavelength, low-energy photons (Fig. 1). This leads to lower photolysis rates on these planets, as there are fewer photons with the requisite energy to dissociate molecules. Figure 1b illustrates this by showing the wavelength-dependent absorption cross section for CH₃SH (Sharpe et al., 2004), along with the incident UV flux from the three different stars. Photolysis of CH₃SH (and the other S_org species) generally occurs at wavelengths <300 nm, where the fluxes from the Sun, AD Leo, and T3100 differ by orders of magnitude. Except below 170 nm, where the AD Leo habitable-zone planet receives the highest relative flux, the UV flux decreases dramatically going from the Sun to AD Leo to T3100. Because CH₃SH photolysis occurs mostly in the 200–300 nm region, its photolysis rate follows this same pattern. The same holds true for other gases, for example, H₂O, whose photolysis creates highly reactive radicals that destroy S_org.

3.1. Production and loss of S_org species

We define a standard Archean model with the general boundary conditions above along with 1 MDF S_org and 1 MDF CH₄. The largest S_org sinks in this simulation were the following reactions:

These reactions outpaced other net S_org sinks by at least an order of magnitude. Thus, the major sink for S_org in our model was reaction with O, and the major by-products were CH₃ and oxidized sulfur species—SO and HSO. The major source of O atoms to the atmosphere was photolysis of major atmospheric components (in an anoxic atmosphere, CO₂, SO₂, and H₂O), and the inventory of O atoms decreased when the flux of UV photons to the atmosphere was diminished.

On the model planet orbiting T3100, the biggest sinks for S_org species were the following reactions:

In the model simulations around these stars, the lack of UV photons entering the atmosphere led to a lack of O radicals in the atmosphere. This caused a slower destruction rate of the S_org gases and shifted the main by-products of S_org photochemistry to carbon monoxide (CO) and reduced sulfur species (S, S₂, and H₂S). The planets orbiting AD Leo were between these two end-member cases for atomic O production. As a result, the by-products of S_org chemistry on planets around M dwarfs were a mix of oxidized and reduced sulfur species.

3.2. Atmospheric profiles

Results from nine photochemical model runs are shown in Figs. 2 and 3. Each figure contains a 3×3 grid of panels with decreasing UV flux (Sun, AD Leo, T3100) from left to right and increasing organic sulfur gases (0 MDF S_org, our control; 1 MDF S_org, the modern-day fluxes; and 10 MDF CH₃SH, corresponding to a biosphere containing mercaptogens) from top to bottom. Figure 2 shows the calculated mixing ratio profiles of the major S_org species along with SO₂ and H₂S, while Fig. 3 shows the calculated vertical profiles of H₂O, CH₄, C₂H₆, H₂, and O₂. The profiles generated with 0 MDF S_org are our control experiments, as this boundary condition is equivalent to assuming no biological S_org production. In these cases, the atmospheric mixing ratios of all S_org gases were extremely low.

FIG. 2. — These nine panels each show model-predicted vertical profiles of the mixing ratios of the organic sulfur species. Panels toward the left are for planets orbiting stars with greater UV radiation, and panels toward the bottom are for planets with higher biological S_org production. The S_org mixing ratios increase with higher ground S_org fluxes (bottom panels) and with lower UV radiation (right panels). Color images available online at www.liebertonline.com/ast

FIG. 3. — These nine panels each show model-predicted vertical profiles of the mixing ratios of the greenhouse gases in our climate and line-by-line radiative transfer models. Panels on the left are for planets orbiting stars with greater UV radiation, and panels on the bottom are for planets with higher biological S_org production. H₂O and CO₂ concentrations are identical in all model runs, while CH₄ concentrations vary only modestly between simulations. Note the increase in C₂H₆ concentrations on planets with higher S_org fluxes or lower UV radiation, or both. Color images available online at www.liebertonline.com/ast

For models with the modern-day S_org flux, near-surface mixing ratios of DMS built up to at least ∼10 ppt (10⁻¹¹) for all three stellar types. These relatively low concentrations are due to higher photolysis rates in the absence of an O₂/O₃ UV shield. For the T3100 model planet, DMDS and CH₃SH peaked above 100 ppb (10⁻⁷). The shapes of the S_org profiles also changed as a function of star type, as the sulfur gases remained well mixed to higher altitudes in the low-UV-flux models, further increasing the total column depths of the S_org species. C₂H₆ concentrations also increased when surface S_org production was included, because of additional production of CH₃ radicals (Fig. 3).

As expected, increasing the CH₃SH flux to 10 MDF while keeping the rest of the S_org gases at 1 MDF (the mercaptogen experiments) resulted in a further increase in all S_org mixing ratios (Fig. 2). This increase was most pronounced in CH₃SH and in DMDS and was greatest on planets receiving relatively low UV radiation. C₂H₆ concentrations also increased with these higher CH₃SH fluxes (Fig. 3), despite the fact that we decreased the CH₄ fluxes in these simulations so that the total flux of CH₃ groups to the atmosphere remained constant.

3.3. Spectra

To illustrate where each of the gases plotted in Figs. 2 and 3 are spectrally active, we present sensitivity spectra of the model planet with 30 MDF CH₃SH orbiting AD Leo (Fig. 4). We generated Fig. 4 by running a full spectral model (shown as a black curve) and then subsequent model runs with one gas removed in each run. These sensitivity spectra are not self-consistent atmospheres; rather, they are tools to determine what gases are causing certain absorption features in the full spectral model. The spectral regions in which a sensitivity spectrum for a particular gas differs from the planet's complete spectrum show where that gas absorbs. For example, the effects of H₂O are clearly seen (difference between black and gray curves) from 5 to 7 μm and longward of 17 μm. Likewise, CO₂ absorption features (difference between black and brown curves) are present from 9 to 11 μm and from 12 to 19 μm, CH₄ absorption is present from 6 to 9 μm, and C₂H₆ has a deep absorption feature from 11 to 13 μm. The distinguishable S_org absorption features include those caused by CH₃SH from 9 to 11 μm and by DMDS from 10 to 11 μm.

FIG. 4. — The top panel shows the absorption cross sections for the gases included in our spectral model. The middle and bottom panels show the simulated spectra for a simulation of a planet with 30 MDF S_org orbiting AD Leo. The black line shows the full model spectrum, including the influence of all the gases in our line-by-line radiative transfer model. The colored lines show model spectra in which one gas is removed from the line-by-line radiative transfer model, with lines of the same color showing the absorption cross-section spectrum for that gas in the top pane. For example, the gray line shows the spectrum with the radiative influence of H₂O removed from the model. The bottom panel shows a zoom-in on the “infrared window” between 8.5 and 11 μm.

Model spectra from 4 to 20 μm are presented at a spectral resolution of R (λ/Δλ)∼50 in Fig. 5. This resolution is consistent with the requirement goal for the Terrestrial Planet Finder Interferometer (TPF-I), a first-generation thermal-IR planet characterization mission (Lawson et al., 2007). For the simulations of a mercaptogen biosphere on a planet with a spectrum of the Sun or AD Leo, the greatest remotely observable difference was the C₂H₆ absorption feature between 11 and 13 μm, the strength of which increases at higher CH₃SH fluxes (30×modern CH₃SH flux). This feature became more prevalent if we increased the flux of the other S_org gases (30×modern S_org flux) or if we decreased the UV radiation reaching the planet (bottom panel). The model simulations with these deeper C₂H₆ features also exhibited enhanced absorption features from 8.5 to 11 μm caused by DMDS.

H₂S fluxes are unlikely to cause false positives. H₂S had a large spectral influence only on planets with extremely large H₂S fluxes (1000×H₂S MDF) orbiting stars with extremely low UV radiation (T3100). Except for these end-member cases, we do not expect H₂S to provide a false negative for the other absorption features discussed here.

4. Discussion

Several trends from our photochemical simulations (Figs. 2 and 3) have implications for the interpretation of future exoplanetary spectra. As the stellar UV flux to the planet decreases, the ground-level mixing ratios and altitudinal extent of S_org species increase. The same effects can also be caused by increases to the S_org surface fluxes. Both trends can be explained by an increase in the ratio of S_org sources to S_org sinks. The main sources of S_org to the atmosphere are the biogenic surface fluxes; an increase in these raises the source/sink ratio. The two main sinks of S_org species are direct photolysis and reaction with radicals such as OH and O that themselves are by-products of photochemical reactions. The decrease in UV radiation slows all photolysis and therefore decreases the sinks for S_org species.

The other robust trend in the photochemical simulations is an increase in C₂H₆ with increasing S_org fluxes and with decreasing UV radiation. Increasing S_org fluxes increases the source of CH₃ radicals that combine to form C₂H₆. Decreases in UV fluxes lead to lower C₂H₆ photolysis rates, lower concentrations of C₂H₆-destroying radicals, and smaller sinks for C₂H₆.

C₂H₆ has not previously been identified as a potential biosignature for anoxic atmospheres, although most concepts for mid-IR exoplanet characterization missions already include plans to detect CH₄ by looking for its absorption feature centered near 7.7 μm (Lawson et al., 2007). According to our model simulations, C₂H₆ detection would require an interferometer with a spectral resolution of λ/Δλ∼20 and a S/N ∼ 15 in the 11–13 μm range to resolve the distinctive band profile for this gas. Such a mission could discriminate at a 3σ level between C₂H₆ produced by the model with the modern-day S_org flux and the model with no S_org flux, for a planet around an M dwarf similar to AD Leo.

C₂H₆ concentrations can be enhanced both by increased S_org concentrations and by increased CH₄. Because CH₄ can have an abiogenic source, CH₄-derived C₂H₆ could be abiogenic in origin. Figure 5 shows low-resolution (R∼50) spectra with high C₂H₆ concentrations arising from either high S_org fluxes or high CH₄ fluxes. Models that have higher S_org fluxes have higher C₂H₆ concentrations and a deeper C₂H₆ absorption feature between 11 and 13 μm. Similarly, models that have higher CH₄ fluxes also have increased C₂H₆ concentrations and more absorption between 11 and 13 μm. However, models that achieve C₂H₆ buildup through increased CH₄ fluxes also exhibit a detectable increase in the CH₄ concentrations in the atmosphere: there was a doubling in the near-surface CH₄ mixing ratios when the CH₄ fluxes were increased to 1.5 MDF, and another doubling when the CH₄ fluxes were increased to 2.0 MDF. These increased CH₄ concentrations caused significantly more absorption between 8 and 9 μm. In other words, changes in the absorption by CH₄ could potentially allow us to discriminate between the spectra with “abiogenic, CH₄-derived C₂H₆” and the spectra with “biogenic, S_org-derived C₂H₆.” Thus, an exoplanet characterization mission that can measure the depths of the CH₄ and C₂H₆ absorption features accurately enough to estimate the C₂H₆/CH₄ ratio may be able to determine whether biological S_org production contributes to the source of C₂H₆.

These above differences in CH₄ absorption depths in biological and abiological model simulations are the result of higher C₂H₆/CH₄ ratios in models with biological S_org fluxes. These fluxes caused an increase in atmospheric CH₃ groups, which in turn increased the atmospheric C₂H₆/CH₄ ratio. Thus, for a given amount of C₂H₆, the CH₄ concentrations were lower in models with higher S_org fluxes. (The converse is also true; for a given CH₄ concentration, models with higher S_org fluxes exhibited higher C₂H₆ concentrations.) This effect could be augmented by inclusion of other biological CH₃X species, such as CH₃Cl, that were not included in these simulations.

In addition to the influence of S_org species on the C₂H₆ feature, several other features were caused directly by the presence of the S_org in the model atmospheres: absorption just shortward of 7 μm by DMS, absorption just longward of 7 μm by DMDS, absorption from 8.5 to 9.5 μm by DMDS, and absorption between 9 and 11 μm by DMDS and CH₃SH. When present, these features created a continuous, but not constant, increase in absorption from 6 μm all the way to the C₂H₆ feature at 11 μm. Thus, they have a significant impact across a wide wavelength range. However, these features only appeared in model simulations with extremely low UV fluxes (the T3100 case) or in simulations with at least 30-fold increases in the flux rate of all S_org gases. On planets around more active stars, these features would only be detectable if the biosphere is much more productive than Earth's biosphere or if the organisms living on the planet have high concentrations of sulfur in their proteins. Even planets with an active mercaptogen community would not produce these features unless that community produces CH₃SH at a rate that is greater than 30 times the modern-day CH₃SH flux from the oceans.

Additional confusion in interpreting potential arises from the influence of surface temperature. Discriminating between planets with absorption by S_org species and planets with lower surface temperatures may prove problematic, as the S_org gases all absorb in the 8–12 μm “atmospheric window” wavelength region. This is a part of the spectrum that some have suggested could be used to discern surface temperatures, because on modern-day Earth that region is the most transparent to the IR radiation emitted by the surface of Earth. However, an increase in greenhouse gases that absorb photons in this region (including S_org species) will increase its opacity, thereby decreasing the effectiveness with which the surface temperature can be ascertained.

The quantitative effect of S_org absorption on inferred planetary temperature is shown by the dashed curves in Fig. 5. Here the model spectra, which are cloud free, have been degraded to the spectral resolution goal for TPF-I and are shown with blackbody spectra at three temperatures: (1) 180 K, the stratospheric temperature in our model (drawn in brown); (2) 278 K, the surface temperature in our model (drawn in goldenrod); and (3) either 276 K (top, middle) or 273 K (bottom), the maximum temperature derived for 1 MDF S_org case within the window region of the model spectrum (drawn in orange). Figure 5 shows that the S_org gas absorption, in addition to weak water vapor absorption, increases the opacity of the atmosphere in the atmospheric window sufficiently that the majority of the radiation sensed comes from higher, colder regions of the planet's troposphere. The discrepancy between actual surface temperature (278 K) and maximum observed temperature is as much as 8 K for the highest S_org fluxes and lowest UV fluxes. This will increase the planet's greenhouse effect but decrease the effectiveness with which the surface temperature can be sensed remotely. This effect is from atmospheric absorption alone and does not account for the atmospheric column-truncating effects of clouds or hazes, which for an unresolved Earth-like planet can further reduce the measured brightness temperature in the window region.

Obtaining the best possible estimates of planetary surface temperatures for extrasolar planets of unknown composition will therefore require sufficient spectral wavelength range and resolution to identify non-Earth-like atmospheric window regions, and good estimates of planetary composition and the presence of cloud or aerosol cover. These measurements, combined with atmospheric modeling, will be crucial for understanding limitations on planetary temperature retrieval from MIR spectra for planets with atmospheric characteristics unlike those of modern Earth. For anoxic atmospheres, it is important to be able to detect S_org absorption features at wavelengths shortward of the window region. Absorption by DMS and DMDS between 6 and 9 μm provides an extra constraint on the abundance of these gases. Similarly, the C₂H₆ feature could be used in conjunction with photochemical models to further constrain the S_org flux rates. The atmospheric S_org inventory could then be input to a climate model to calculate self-consistent surface temperatures and spectra. A fairly comprehensive characterization of an anoxic atmosphere could therefore be achieved with spectra from 6 to 13 μm (and preferably down to 5 μm and out to 20 μm to help constrain water abundances) at a spectral resolution of at least 20 and a S/N greater than 15. These baseline parameters are consistent with the current requirement goals for the TPF-I mission concept.

5. Conclusions

In this paper, we have shown that an anoxic biosphere could be detected over interstellar distances by searching for organic S species produced by biology. On planets orbiting Sun-type stars, S_org fluxes at 30 times modern-day levels could be detected in the form of elevated C₂H₆/CH₄ ratios that are a photochemical by-product of S_org gases. On planets around M dwarfs such as AD Leo, detection of heightened C₂H₆/CH₄ ratios is possible at present-day S_org fluxes. Features caused directly by S_org gases may be observable on planets that have much higher S_org fluxes or on planets orbiting M dwarfs that exhibit low amounts of stellar activity, or both. An important caveat to this work is that aerosols, including water clouds, hydrocarbon aerosols, sulfate aerosols, and S₈ particles, were not considered in the spectral portion of this study but may impact the ability to detect these species.

The detection of any of these features will require an instrument with spectral resolution R>20, broad coverage of the IR spectrum (6–14 μm), and low total noise levels (S/N>15 or noise < 1 W/m²/μm). Current expected performance levels for TPF-I meet these requirements (Lawson et al., 2007). The use of models to interpret the spectra will also be required in order to separate the effects of surface temperature, organic sulfur gases, and other atmospheric constituents on the planetary spectrum.

Despite the difficulties involved, the benefits offered by such a search are considerable. By including organic sulfur species in our repertoire of remotely detectable biosignatures, the detection of life on some planets may also come with rudimentary lessons on the composition of that planet's biosphere. Thus, this work supports exoplanet characterization missions with a wavelength range and spectral resolution sufficient to detect C₂H₆ and the organic sulfur gases discussed above.

Acknowledgments

The quality of this manuscript was improved by the useful critiques of three reviewers. This work was performed as part of the NASA Astrobiology Institute's Virtual Planetary Laboratory, supported by the National Aeronautics and Space Administration through the NASA Astrobiology Institute under solicitation No. NNH05ZDA001C. Additional support was provided by the Pennsylvania State Astrobiology Research Center, supported by the National Aeronautics and Space Administration through the NASA Astrobiology Institute under grant number NNA09DA76A. S.D.D.-G. acknowledges additional support from the Astrobiology NSF IGERT grant at the University of Washington. M.C. and S.D.D.-G. acknowledge additional support from the NASA Postdoctoral Program. J.F.K. acknowledges additional support from NASA's Exobiology and Evolutionary Biology Program.

Abbreviations

DMS, dimethyl sulfide; DMDS, dimethyl disulfide; MDF, modern-day flux; S/N, signal-to-noise ratio; TPF-I, Terrestrial Planet Finder Interferometer.

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PERMALINK

Using Biogenic Sulfur Gases as Remotely Detectable Biosignatures on Anoxic Planets

Shawn D Domagal-Goldman

Victoria S Meadows

Mark W Claire

James F Kasting

Abstract

1. Introduction

2. Methods

2.1. Photochemical code

Table 1.

2.2. Boundary conditions

Table 2.

FIG. 1.

2.3. Radiative transfer code

3. Results

3.1. Production and loss of S_org species

3.2. Atmospheric profiles

FIG. 2.

FIG. 3.

3.3. Spectra

FIG. 4.

FIG. 5.

4. Discussion

5. Conclusions

Acknowledgments

Abbreviations

References

ACTIONS

PERMALINK

RESOURCES

Cite

Add to Collections

PERMALINK

Using Biogenic Sulfur Gases as Remotely Detectable Biosignatures on Anoxic Planets

Shawn D Domagal-Goldman

Victoria S Meadows

Mark W Claire

James F Kasting

Abstract

1. Introduction

2. Methods

2.1. Photochemical code

Table 1.

2.2. Boundary conditions

Table 2.

FIG. 1.

2.3. Radiative transfer code

3. Results

3.1. Production and loss of Sorg species

3.2. Atmospheric profiles

FIG. 2.

FIG. 3.

3.3. Spectra

FIG. 4.

FIG. 5.

4. Discussion

5. Conclusions

Acknowledgments

Abbreviations

References

ACTIONS

PERMALINK

RESOURCES

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3.1. Production and loss of S_org species