BIOSIGNATURE GASES IN H₂-DOMINATED ATMOSPHERES ON ROCKY EXOPLANETS

Photochemistry Model. The focus on chemistry is critical for biosignature gases because sinks control a biosignature gas lifetime and hence the gas' potential to accumulate in the planetary atmosphere. A model for atmospheric chemistry is required to connect the amount of biosignature gas in the atmosphere (as required for detection) to the biosignature source flux at the planetary surface.

Our photochemical model is presented in Hu et al. (2012). The photochemical code computes the steady-state chemical composition of an exoplanetary atmosphere. The system can be described by a set of time-dependent continuity equations, one equation for each species at each altitude. Each equation describes chemical production, chemical loss, eddy diffusion and molecular diffusion (contributing to production or loss), sedimentation (for aerosols only), emission and dry deposition at the lower boundary, and diffusion-limited atmospheric escape for light species at the upper boundary. The code includes 111 species, 824 chemical reactions, and 71 photochemical reactions. Starting from an initial state, the system is numerically evolved to the steady state in which the number densities no longer change.

The generic model computes chemical and photochemical reactions among all O, H, N, C, and S species and the formation of sulfur and sulfate aerosols. The numerical code is designed to have the flexibility of choosing a subset of species and reactions in the computation. The code therefore has the ability to treat both oxidized and reduced conditions by allowing the selection of "fast species." For the chemical and photochemical reactions, we use the most up-to-date reaction rate data from both the NIST database (http://kinetics.nist.gov) and the JPL publication (Sander et al. 2011). Ultraviolet (UV) and visible radiation in the atmosphere is computed by the δ-Eddington two-stream method with molecular absorption, Rayleigh scattering, and aerosol Mie scattering contributing to the optical depth. The model was developed from the ground-up and has been tested and validated by reproducing the atmospheric composition of Earth and Mars (Hu et al. 2012; Hu 2013).

For biosignature gases that are minor chemical perturbers in the atmosphere, the biosignature lifetime can be estimated based on the abundance of the major chemical sink. For this paper, the values of NH₃ and N₂O surface source fluxes are calculated from the full photochemistry model, whereas the calculations of surface source fluxes for other biosignature gases are simplified estimates. One more point to note is that photochemistry is relatively high in the atmosphere, typically above mbar levels, where stellar UV radiation can penetrate from above.

Temperature–Pressure Profile. The precise temperature-pressure structure of the atmosphere is less important than photochemistry for a first-order description of biosignatures in H₂-rich atmospheres. The reason is that most biosignature gases of interest have sources and sinks that are not significantly affected by minor deviations in the temperature pressure profile. Moreover, the biosignature gases themselves are secondary players in governing the heating structure of the atmosphere.

We therefore justify using the photochemistry model in a stand-alone mode, with a pre-calculated temperature–pressure profile. The calculated temperature–pressure profile is approximate and is one that assumes a surface temperature (i.e., 290 K), an appropriate adiabatic lapse rate for H₂-rich compositions, and a constant temperature above the convective layer of the atmosphere. Such assumed temperature profiles are consistent with greenhouse warming in the troposphere and a lack of UV absorbers in the stratosphere. The semi-major axis of the planet is then derived based on the assumed temperature profile by balancing the energy flux of incoming stellar radiation and outgoing planetary thermal emission. The details of this procedure are described in Hu et al. (2012).

Synthetic Spectra. To generate exoplanet transmission and thermal emission spectra, we use a line-by-line radiative transfer code (Seager et al. 2000; Madhusudhan & Seager 2009; Hu et al. 2012). Opacities are based on molecular absorption with cross sections computed based on data from the HITRAN 2008 database (Rothman et al. 2009), molecular collision-induced absorption when necessary (e.g., Borysow 2002), Rayleigh scattering, and aerosol extinction computed based on Mie theory. The atmospheric transmission is computed for each wavelength by integrating the optical depth along the limb path (as outlined in, e.g., Seager & Sasselov 2000; Miller-Ricci et al. 2009). The planetary thermal emission is computed by integrating the radiative transfer equation without scattering for each wavelength (e.g., Seager 2010).

We consider clouds in the emergent spectra for thermal emission by considering 50% cloud coverage by averaging cloudy and cloud-free spectra. We omit clouds for the transmission spectra model because the clouds are at low altitudes whereas the spectral features form at high altitudes.

A biomass model estimate has been developed by Seager et al. (2013) that ties biomass surface density to a given biosignature gas surface source flux. The motivating rationale is that with a biomass estimate, biosignature gas source fluxes can be free parameters in model predictions and therefore provide a physical plausibility check in terms of reasonable biomass. The approach aims to enable consideration of a wide variety of both gas species and their atmospheric concentration in biosignature model predictions. The biomass model estimates are valid to one or two orders of magnitude. We provide a summary of the biomass model here with the full details available in Seager et al. (2013).

The biomass model is used in the following algorithm. First, we calculate the amount of biosignature gas required to be present at "detectable" levels in an exoplanet atmosphere from a theoretical spectrum (we define a detection metric in Section 2.3). Second, we determine the gas source flux necessary to produce the atmospheric biosignature gas in the required atmospheric concentration. The biosignature gas atmospheric concentration is a function not only of the gas surface source flux, but also of other atmospheric and surface sources and sinks. Third, we estimate the biomass that could produce the necessary biosignature gas source flux. Fourth, we consider whether the estimated biomass surface density is physically plausible, by comparison with maximum terrestrial biomass surface density values and total plausible surface biofluxes.

Based on life on Earth, a summary overview is that a biomass surface density of 10 g m⁻² is sensible, 100 g m⁻² is plausible, and 5000 g m⁻² is possible. In real situations, the total biomass is nearly always limited by energy, bulk nutrients (carbon, nitrogen), trace nutrients (iron, etc.), or all three. Regarding global surface biofluxes, we provide values and references where needed in our results and discussion.

The biomass model estimates are tied to the type of biosignature gas and so we briefly summarize our biosignature classification scheme before discussing each biomass model estimate.

Type I biosignature gases are generated as by-product gases from microbial energy extraction. For example, on Earth, many microbes extract energy from chemical energy gradients using the abundant atmospheric O₂ for aerobic oxidation:

$$\begin{equation} {\rm X + O_2 \rightarrow oxidized X }. \end{equation} \tag{ 2 }$$

H₂O is generated from H₂, CO₂ from organics, SO₂ or SO$_4^{2-}$ from H₂S, rust from iron sulfide (FeS), NO$_2^{-}$ and NO$_3^{-}$ from NH₃, etc.

On an exoplanet with an H₂-rich atmosphere, the abundant reductant would now be atmospheric H₂ such that

$$\begin{equation} {\rm H_2 + X \rightarrow reduced X }. \end{equation} \tag{ 3 }$$

The oxidant must come from the interior.

In other words, for chemical potential energy gradients to exist on a planet with an H₂-rich atmosphere, the planetary crust must (in part) be oxidized in order to enable a redox couple with the reduced atmosphere. The by-product is always a reduced gas, because in a reducing environment H₂-rich compounds are the available reductants. To be more specific, oxidants would include gases such as CO₂ and SO₂.

The Type I biosignature gas biomass model is based on thermodynamics and is derived from conservation of energy and discussed in detail in Seager et al. (2013). The biomass model estimate is

$$\begin{equation} \Sigma _B \simeq \Delta G \left[ \frac{F_{\rm source}}{ P_{m_e}} \right]. \end{equation} \tag{ 4 }$$

Here, Σ_B is the biomass surface density in g m⁻² and ΔG is the Gibbs free energy of the chemical redox reaction from which energy is extracted (i.e., Equation (2)). ΔG depends on the standard free energy of reaction (ΔG₀) and the concentration of the reactants and products. Reactant and product concentrations can include ocean pH (concentration of H⁺) in reactions that generate or consume protons. ΔG₀ values are taken from Amend & Shock (2001).

The term $P_{m_{e}}$ is an empirically determined microbial maintenance energy consumption rate that is the minimum amount of energy an organism needs per unit time to survive in an active state (i.e., a state in which the organism is ready to grow). An empirical relation has been identified by Tijhuis et al. (1993) that follows the Arrhenius law

$$\begin{equation} P_{m_e} = A \exp \left[ \frac{-E_A}{RT} \right]. \end{equation} \tag{ 5 }$$

Here, E_A = 6.94 × 10⁴ kJ mol⁻¹ is the activation energy, R = 8.314 kJ mol⁻¹ K⁻¹ is the universal gas constant, and T in units of K is the temperature. The constant A is 4.3 × 10⁷ kJ g⁻¹ s⁻¹ for aerobic growth and 2.5 × 10⁷ kJ g⁻¹ s⁻¹ for anaerobic growth (Tijhuis et al. 1993). Here, per g refers to per g of the wet weight of the organism. $P_{m_e}$ is in units of kJ g⁻¹ s⁻¹.

The free parameter in this biomass model estimate (Equation (4)) is the biosignature gas source flux F_source (in units of mole m⁻² s⁻¹). F_source is the flux of the metabolic by-product and is also the surface bioflux required to generate a given biosignature gas concentration in the atmosphere.

Type II biosignature gases are by-product gases produced by the metabolic reactions for biomass building and require energy. On Earth, these are reactions that capture environmental carbon (and, to a lesser extent, other elements) in biomass. Type II biosignature reactions are energy-consuming and on Earth the energy comes from sunlight via photosynthesis.

There is no useful biomass model for Type II biosignature gases because once the biomass is built a Type II biosignature gas is no longer generated.

Type III biosignature gases are produced by life but not as by-products of their central chemical functions. Type III biosignature gases appear to be particular to specific species or groups of organisms and require energy for their production. Because the chemical nature and amount released for Type III biosignature gases are not linked to the local chemistry and thermodynamics, the Type III biosignature gas biomass model is an estimate based on lab culture production rates.

We estimate the biomass surface density by taking the biosignature gas source flux F_source (in units of mole m⁻² s⁻¹) divided by the mean gas production rate in the lab R_lab (in units of mole g⁻¹ s⁻¹):

$$\begin{equation} \Sigma _B \simeq \frac{F_{\rm source}}{R_{\rm lab}}. \end{equation} \tag{ 6 }$$

We take the maximum observed for the Type III R_lab rates from different studies (Seager et al. 2013). The caveat of the Type III biomass estimate explicitly assumes that the range of R for life on exoplanets is similar to that for life in Earth's lab environment. Nonetheless, we have showed that the Type III biomass model is valid to one or two orders of magnitude, based on Earth's values. The goal, again, is to use the biomass estimate to argue for or against the plausibility of a biosignature gas based on Earth's biomass surface density values, not to make any predictions of quantitative values.

Bioindicators are defined as the end product of the chemical reactions of a biosignature gas.

Model caveats are related to the order-of-magnitude nature of the biomass estimates, the possible terracentricity of the biomass model estimates, and the lack of ecosystem context (see Seager et al. 2013, Sections 6.1– 6.3 for a detailed discussion). Here, we provide a summary overview.

The order-of-magnitude nature of the Type I biomass estimate derives from the dependency of the estimate on $P_{m_e}$, itself very sensitive to temperature. The possible terracentricity of our estimates is related to the use of $P_{m_e}$, which is derived from observations of terrestrial microorganisms, but we argue that the dependency is largely based on thermodynamics (Seager et al. 2013). The order-of-magnitude nature of the Type III biosignatures derives from the reliance on laboratory rates for microbial production rates; this is also possibly a terracentricity issue.

The lack of ecosystem context is a major limitation for the biomass estimate. An ecosystem contains not only the producers (i.e., the biomass estimate derived ultimately from the bioflux F_source) but also the consumers, whereas the biomass model estimate considers only the producers. In this sense, the biomass estimate in Equation (4) is a minimum. We can fairly say that in the case of a very small or very large biomass estimate, the assessment of biosignature gas plausibility is valid: a small biomass estimate gives room for consumers even as a minimum biomass and a large biomass estimate as a minimum will remain large regardless of the consumers. For the intermediate case where a large but not unreasonable biomass is needed to generate a detectable biosignature, the decision of whether the gas is a plausible biosignature is more complicated and will depend on the planetary context: geochemistry, surface conditions, atmospheric composition, and other factors.

Again, we do not argue that the biosignature biomass model estimates are an accurate prediction of an extraterrestrial ecology. Rather, we emphasize that the goal of the biomass model estimates is the order-of-magnitude nature for a first-order assessment of the plausibility of a given biosignature gas candidate.

We now describe our metric for a "detection" that leads to a required biosignature gas concentration. For now, detection has to be a theoretical exercise using synthetic data. We determine the required biosignature gas concentration based on a spectral feature detection with a signal-to-noise ratio (S/N) = 10. Specifically, we describe the S/N of the spectral feature as the difference between the flux in the absorption feature and the flux in the surrounding continuum (on either side of the feature), taking into account the uncertainties on the data:

$$\begin{equation} {{\rm S}/{\rm N}} = \frac{\left| F_{{\rm out}} - F_{{\rm in}} \right|}{\sqrt{\sigma ^2_{F_{{\rm out}}} + \sigma ^2_{F_{{\rm in}}}}}, \end{equation} \tag{ 7 }$$

where $F_{{\rm in}} \pm \sigma _{F_{{\rm in}}}$ is the flux density inside the absorption feature, $F_{{\rm out}} \pm \sigma _{F_{{\rm out}}}$ is the flux density in the surrounding continuum, and σ is the uncertainty on the measurement.

The uncertainties of the in-feature flux and continuum flux are calculated for limiting scenarios. For thermal emission, we consider a futuristic space telescope able to block out the light of the host star. The uncertainties of the in-feature flux and continuum flux are calculated for a limiting scenario: a 1.75 times Earth-sized planet orbiting a star⁶ at 10 pc observed (via direct imaging) with a 6 m diameter telescope mirror operating within 50% of the shot noise limit and a quantum efficiency of 20%. The integration time is assumed to be 20 hr. We note that collecting area, observational integration time, and source distance are interchangeable depending on the time-dependent observational systematics. This telescope scenario is based on a TPF-I type telescope (Lawson et al. 2008).

For transit transmission spectra, we use the same equation as above but with the denominator replaced by the noise in the stellar flux (F_*), as in $\sqrt{{A^A}\smash{{{(4\sigma _F^2)}}}}$, because transmission observations measure the difference between the in-transit stellar flux and out-of-transit stellar flux. For transmission spectra, we consider a 6.5 m space telescope, having a quantum efficiency of 25% observing with a 50% photon noise limit, with integration time of 60 hr in-transit and 60 hr out of transit (assuming observing of multiple transits). Again, we note that collecting area, observational integration time, and source distance are interchangeable depending on the time-dependent observational systematics. This scenario is based on the JWST.

We start by focusing on the Type I biosignature gases, gases generated by reactions that extract energy from external, environmental redox gradients. The most likely Type I metabolic product in an H₂-rich atmosphere would be that in which non-hydrogen elements are in their most hydrogenated form.⁷ In a reducing environment, life captures chemical energy by reducing environmental chemicals. In the presence of excess hydrogen, the most energy that life could extract from chemical potential energy gradients would be from converting elements from relatively oxidized compounds to their fully reduced form. An additional reason for focusing on Type I biosignature gas products that are in their most hydrogenated form is that they are likely long lived in an H₂-dominated atmosphere because molecules in their most hydrogenated form cannot undergo any further reactions with H.

4.1.1. Type I Biosignature Gas Overview

4.1.2. NH₃ as the Strongest Candidate Biosignature Gas in an H₂ Atmosphere

NH₃ is the strongest candidate biosignature gas in a thin, H₂ atmosphere because, like O₂ in Earth's atmosphere, there is no plausible geological or photochemical mechanism for producing high concentrations on rocky planets with thin atmospheres (but cf. the false positive discussion below). NH₃ is readily photolyzed in the upper atmosphere to yield N₂ and is thermally broken down in volcanic gases at high temperatures. The triple bond of N₂ makes it extremely kinetically stable and so any N in the atmosphere ends up being trapped as N₂.

We have therefore proposed NH₃ as a biosignature gas in an H₂-rich atmosphere (Seager et al. 2013). NH₃ is a good biosignature gas candidate for any thin H₂-rich exoplanet atmosphere because of its short lifetime and lack of geological production sources. NH₃ as a biosignature gas is a new idea and one that is specific to a non-Earth-like planet. On Earth, NH₃ is not a useful biosignature gas because, as a highly valuable molecule for life that is produced in only small quantities, it is rapidly depleted by life and is unable to accumulate in the atmosphere. NH₃ is also a very poor biosignature gas on Earth because it is very soluble, so the trace amounts produced will stay dissolved in water and will not escape to the atmosphere.

The summary of the biosignature gas idea is that NH₃ would be produced from hydrogen and nitrogen, in an atmosphere rich in both:

$$\begin{equation} {\rm 3H_2 + N_2 \rightarrow 2NH_3.} \end{equation} \tag{ 18 }$$

This reaction is exothermic and could be used to capture energy. The industrial version of this reaction is called the Haber process for ammonia production at high temperatures; hence, we call such a planet a Cold Haber World. We proposed that in an H₂-rich atmosphere, life could find a way to catalyze the breaking of the N₂ triple bond and the H₂ bond to produce NH₃ and capture the energy released. In contrast, life on Earth solely fixes nitrogen in an energy-requiring process. Energy capture would yield an excess of NH₃ over that needed by life to build biomass, so the excess would accumulate in the atmosphere. Is a Cold Haber World possible? We believe yes, based on synthetic chemistry on Earth that can catalyze the breakage of each of the H₂ (Nishibayashi et al. 1998) and N₂ bonds (Yandulov & Schrock 2003; Schrock 2011) at Earth's surface pressure and temperature; what is not yet known is a catalytic system that can break both at once.

We showed in Seager et al. (2013) that for an Earth-sized, Earth-mass planet with a 1 bar atmosphere of 75% N₂ by volume and 25% H₂ by volume (including carbon species via a CO₂ emission flux), a potentially detectable NH₃ atmosphere concentration of 0.1 ppm is sustainable by a very reasonable biomass surface density of 9 × 10⁻² g m⁻². This modest surface density corresponds to a layer less than one bacterial cell thick. For comparison, the phytoplankton that are the major contributor to Earth's oxygen atmosphere are present in Earth's oceans at around 10 g m⁻². For interest, we note that standard printer paper is between 80 and 100 g m⁻².

For an H₂-dominated atmosphere with 90% H₂ and 10% N₂ on a planet with 10 M_⊕ and 1.75 R_⊕ orbiting a Sun-like star, but all other parameters the same as the above, the viability of NH₃ as a biosignature gas in a thermal emission spectrum still holds based on a physically reasonable biomass surface density. We now describe the estimate for the biomass surface density using the Type I biomass equation (Equation (4)). We use the NH₃ source flux of 2.4 × 10¹⁵ molecule m⁻² s⁻¹ (see Table 2). To compute ΔG, we used T = 290 K and reactant and product concentrations at the surface in terms of partial pressures of N₂ = 0.1, H₂ = 0.9, and NH₃ = 1.4 × 10⁻⁷, giving ΔG = 85.6 kJ mole⁻¹. With $P_{m_e} = 7.0 \times 10^{-6}$ kJ g⁻¹ s⁻¹, we find a biomass surface density of 4.9 × 10⁻² g m⁻². Based on this reasonable biomass surface density, we therefore consider the NH₃ production flux to be viable in our Haber World scenario. The global annual biogenic NH₃ surface emission in the Haber World would be about 1100 Tg yr⁻¹. This is much higher than Earth's natural NH₃ emission at 10 Tg yr⁻¹ (Seinfeld & Pandis 2000). Comparing NH₃ production on the Haber World and on Earth, however, is not valid. We are postulating that production of NH₃ on the Haber World is a major source of metabolic energy for life. A better emission rate comparison is to compare the biosignature gas O₂ from Earth's principal energy metabolism, photosynthesis. The Earth's global oxygen flux is 200 times larger than the Haber World's NH₃ surface emission, at about 2 × 10⁵ Tg yr⁻¹ (Friend et al. 2009).

Turning to a weakly active M5V dwarf star, for the same fiducial planet, the NH₃ surface flux required to sustain a detectable level of atmospheric NH₃ in a thermal emission spectrum is 5.1 × 10¹⁴ molecule m⁻² s⁻¹ (see Table 2). This value is about five times lower than the Sun-like star example above and therefore converts into a biomass estimate about five times smaller than the Sun-like star example above, or about 1 × 10⁻² g m⁻², due to the linear scaling of the problem. For a transmission spectrum measurement for the same planet orbiting the same weakly active M dwarf star, the optimal wavelength range for detection is 2.8–3.2 μm, the required concentration is 11 ppm, and the required surface source flux is 5.5 × 10¹⁶ molecule m⁻² s⁻¹, resulting in a surface biomass of 1 g m². For this particular example, NH₃ in transmission versus thermal emission, it is more difficult to detect NH₃ and hence a higher biomass is actually required for the same hypothetical type of life.

We emphasize that the NH₃ biosignature gas concept is not changed for a planet with a massive (yet still "thin") atmosphere with high surface pressure. As long as the surface conditions are suitable for liquid water, NH₃ will not be created by uncatalyzed chemical reactions.

NH₃ is not immune to false positives. Although a rocky planet with a thin H₂-dominated atmosphere is unlikely to have an NH₃ false positive, the challenge is in identifying the planetary (and stellar) characteristics. We describe three scenarios that could lead to the nonbiological production of NH₃.

A rocky world with a hot surface of ∼820 K could generate NH₃ by the conventional Haber process if there is surface iron. Such a hot surface temperature could presumably be ruled out from other observations.

A second scenario where NH₃ is naturally occurring is in the atmospheres of gas giant planets or the so-called mini-Neptunes. The deep atmosphere may reach conditions where NH₃ can be formed kinetically at the extremely high pressures necessary for NH₃ formation to be possible thermodynamically. On Jupiter, for example, the H₂ + N₂ → NH₃ reaction becomes significant in comparison with vertical transport at about 1500 K and 1400 bar (Prinn & Olaguer 1981). The only way we can discriminate between planets with a massive envelope and a rocky planet with a thin atmosphere where the pressures for the thermodynamic formation of NH₃ are not reached is with high-resolution spectra to assess the surface pressure (Benneke & Seager 2012, 2013).

A third scenario for an NH₃ false positive is for planets with outgassed NH₃ during evolution. The importance of ammonia for the atmospheric evolution of Titan relates to primordial ammonia that accreted with the ices of the moon and has not subsequently been broken down either by internal heat (likely on a rocky planet) or by external UV photolysis (which will rapidly break down any NH3 in the atmosphere; Shin et al. 2012). In this case, ammonia is therefore present as ice in the interior. This would be a challenging case to ascertain and illustrates how an assignment of any gas as a biosignature gas candidate has to be given a detailed probabilistic assessment based on what we know about the relevant planet.

For any case, for a quiescent M star with no chromospheric UV emission—hence a planet with little to no destructive UV flux—NH₃ can easily accumulate in the planet atmosphere and act as a significant false positive. NH₃ is destroyed by photolysis and is very sensitive to the amount of UV radiation.

4.1.3. CH₄ and H₂S as Unlikely Biosignature Gases

CH₄ has been described at length as a possible biosignature gas on early Earth and on exoplanets with oxidized atmospheres (e.g., Hitchcock & Lovelock 1967; Des Marais et al. 2002). This is despite the risk of a geologically derived false positive, because it is believed that in an oxidized environment geological production of CH₄ will be small, and so if enough CH₄ is produced it may be attributed to life. This is the case on Earth where at least 99% of the atmospheric CH₄ derives directly from life or from industrial destruction of fossil hydrocarbons formed from past life (Wang et al. 2004). However, the 1775 ppb concentration of CH₄ in Earth's atmosphere (Solomon et al. 2007) is not enough to be detected remotely with envisioned space telescope capabilities.

CH₄ is a poor biosignature gas in an H₂-rich atmosphere because it is both produced volcancially and is an end product of CO₂ photochemistry in the atmosphere. Terrestrial volcanic emission rates of CH₄ and CO₂ would lead to substantial buildup of CH₄ in H₂-dominated atmospheres. Even small amounts of outgassed CO₂ will lead to an accumulation of CH₄ in the atmosphere because CH₄ has a very long lifetime in an H₂-rich atmosphere and CH₄ would be produced by

$$\begin{equation} {\rm 4H_2 + CO_2 \rightarrow CH_4 + 2H_2O}. \end{equation} \tag{ 19 }$$

More specifically, considering Earth's volcanic emission rates of CH₄ and CO₂, and with deposition velocities of 10⁻⁶ m s⁻¹ for CO₂ and 0 m⁻² s⁻¹ for CH₄, CH₄ will accumulate up to ∼10 ppm in a 1 bar 90% H₂ and 10% N₂ atmosphere with a temperature profile similar to that of the Earth (Hu et al. 2012). Even in the case of no surface CH₄ emission, CO₂ emission into the same atmosphere would lead to the atmospheric production and accumulation of CH₄ up to 5 ppm. This example is intended to show that the false positive risk of CH₄ is so high in an H₂-dominated atmosphere as to make CH₄ an implausible biosignature gas.

H₂S is even more unfavorable than CH₄ as a biosignature gas in an H₂-rich atmosphere because of the same geological false positive issues as with carbon gases. An added problem is the generation of aerosols that may blanket any spectral features and the fact that the H₂S spectral features are heavily contaminated by atmospheric water vapor, making them potentially difficult to detect (Hu et al. 2013).

Type II biosignature gases are those produced by metabolic reactions for biomass building. Biomass building on Earth primarily occurs by photosynthesis, which has the dual goal of harvesting light energy to use for metabolism and also capturing carbon for biomass building.

We have not identified any useful biosignature gases of Type II in an H₂-rich, 1 bar atmosphere. Photosynthesis in a reduced environment such as an H₂-dominated atmosphere would generate reduced by-product gases, which are not useful as biosignature gases because those species are already expected to be present in their most reduced forms in the H₂-dominated atmosphere.

The concept of photosynthesis on a planet with an H₂-dominated atmosphere is nonetheless worth some discussion,⁸ starting with a brief review of photosynthesis in the familiar Earth environment. Photosynthesis must convert carbon from its environmental form, which is the form that is most thermodynamically stable at surface temperatures and pressures, into biomass. The biomass is of an intermediate redox state (Bains & Seager 2012). The key point therefore is that in an oxidized environment like that of the Earth, photosynthesis must reduce oxidized carbon (CO₂) and will generate an oxidized by-product. On Earth, environmental carbon is captured in photosynthesis, producing O₂ as a by-product:

$$\begin{equation} {\rm H_2O + CO_2 \rightarrow CH_2O + O_2}. \end{equation} \tag{ 20 }$$

Here, CH₂O represents biomass.

Photosynthesis, by definition, will have the same goal in an H₂-rich environment as in an oxidized environment: to harvest light energy and build carbon-based biomass. Because CH₄ is the most thermodynamically stable gaseous form of carbon in this environment, photosynthesis would oxidize the carbon in CH₄ and produce a reduced by-product. The lowest energy route is to directly split the CH₄ as

$$\begin{equation} {\rm CH_4 + H_2O + X \rightarrow CH_2O + XH}, \end{equation} \tag{ 21 }$$

where X is an atom that is oxidized in the environment and has been reduced to XH, consuming energy in the process, and CH₂O again represents biomass. (We note that the oxidation state of the oxygen is not changed in this process, unlike oxygenic photosynthesis on Earth, so formally this is not splitting water even though water is involved.)

The null result for photosynthetic biosignatures in an H₂-dominated atmosphere is based on the point that most non-metals (C, O, S, the halogens) are likely to be in their most reduced state already on the surface of this world and so cannot play the role of X in the photosynthesis process described above.

One exception might have been hydrogen, which is oxidized in water and methane and so a possible photosynthetic reaction is

$$\begin{equation} {\rm CH_4 + H_2O \rightarrow CH_2O + 2H_2}, \end{equation} \tag{ 22 }$$

but again H₂ is not a useful biosignature gas because it is already present in the H₂-dominated atmosphere.

For completeness, we describe some other unlikely but interesting possibilities for X and XH. Silicon, phosphorus, and boron are likely to be present as the oxidized minerals silicates, phosphates, and borates, respectively, but using these as a sink for the electrons in photosynthesis, for example in the reaction with silica to generate silane,

$$\begin{equation} {\rm CH_4 + \frac{1}{2} SiO_2 \rightarrow CH_2O + \frac{1}{2} SiH_4}, \end{equation} \tag{ 23 }$$

requires more energy than the reaction in Equation (22) under a range of conditions and so would represent a very inefficient way of generating biomass.

Reduction of a metal with a positive electrochemical potential would be more energetically efficient, as for example in the reduction of copper oxide to copper:

$$\begin{equation} {\rm CH_4 + 2Cu(O) \rightarrow CH_2O + 2Cu + H_2O}, \end{equation} \tag{ 24 }$$

but this reduction produces no volatile product and is dependent on a supply of oxidized metal. (There are clear parallels with anoxygenic photosynthesis on Earth for this type of reaction.) In contrast, the reaction in Equation (22) is limited only by the supply of methane, as life in water is not limited by the chemical availability of water. In summary, photosynthesis in the reducing environment will either generate H₂, which will not be detectable in a hydrogen-dominated atmosphere, or will produce non-volatile products, i.e., products not in gas form, which, by definition, will not be detectable as atmospheric gases.

Life produces many molecules for reasons that are not related to the generation of energy, which we refer to as Type III biosignature gases. The gases are produced for reasons such as stress, signaling, and other physiological functions, and some of these have already been discussed quantitatively in detail as biosignature gases in oxidized atmospheres, (e.g., CH₃Cl (Segura et al. 2005) and DMS and other sulfur compounds (Domagal-Goldman et al. 2011)).

The fate of Type III biosignature gas molecules depends on the level of relevant reactive species in the atmosphere and hence on stellar UV flux. In low-UV environments, some Type III gases can accumulate to detectable levels. In the relatively high-UV environments of Sun-like stars, many Type III gases could be rapidly driven to their most hydrogenated form in an H₂-rich atmosphere. In some cases, they will not accumulate to detectable levels unless we assume unrealistic production rates. In these extreme cases in a high-UV environment, we would only be able to infer the presence of the biosignature gas by detecting the end product of photochemical attack, which we call a bioindicator. Only in a few cases might bioindicators be useful, because many are not spectroscopically active (and hence not detectable) and others are indistinguishable from geological cases as well (e.g., DMS will end up as CH₄ and H₂S; N₂O will end up as N₂ and H₂O.)

We now show that Type III biosignature gas survival and hence plausibility depends highly on the UV flux level of the host star. We consider the three fiducial stellar types that differ in UV radiation levels: a Sun-like star, a weakly active M5V dwarf star, and a quiescent M5V dwarf star (Figure 1). We consider the same model planet as above, a 10 M_⊕, 1.75 R_⊕ planet with an atmosphere with 90% H₂ and 10% N₂ by volume, with a surface temperature of 290 K and a surface pressure of 1 bar. Results for the cases we modeled are listed for thermal emission spectra in Table 2 and for transmission spectra in Table 3.

Our first example of a Type III biosignature gas is methyl chloride (CH₃Cl). CH₃Cl is produced in trace amounts by many microorganisms on Earth. The detectability of CH₃Cl in Earth-like atmospheres in the low-UV environment of UV-quiet M stars has already been studied by Segura et al. (2005) and later as a potential biosignature gas in more generalized oxidized atmospheres by Seager et al. (2013).

Here, for the first time, we study CH₃Cl as a potential biosignature gas in a thin H₂-rich atmosphere. For this, we go beyond previous work not only by considering an H₂ atmosphere but also by using our biomass estimate framework so as not to be constrained by terrestrial bioflux production rates. We now show why CH₃Cl is a potential biosignature gas in H₂-rich atmospheres in low-UV environments—because the amount of biomass needed to generate a detectable concentration of CH₃Cl is physically plausible. We use our biomass estimate framework (Section 2.2 and Seager et al. 2013).

Considering the thermal emission spectrum for our fiducial planet with a 1 bar atmosphere of 90% H₂ and 10% N₂, a spectral signature of 9 ppm is required for spectral detection using our detection metric. This statement is for a spectral band feature in absorption at 13.0–14.2 μm (see Figure 4); this is the band accessible in an H₂ atmosphere, weaker than the 6.6–7.6 μm band that would be masked by H₂–H₂ collision-induced absorption. In order to sustain an atmospheric concentration of 9 ppm of CH₃Cl on our model planet in the habitable zone for a Sun-like star, a weakly active M5V dwarf star, and a UV-quiet M5V dwarf star, the surface bioflux production rate would need to be 1.0 × 10¹⁷ molecule m⁻² s⁻¹ (1.7 × 10⁻⁷ mole m⁻² s⁻¹), 2.9 × 10¹⁵ molecule m⁻² s⁻¹ (4.8 × 10⁻⁹ mole m⁻² s⁻¹), and 4.7 × 10¹¹ molecule m⁻² s⁻¹ (7.8 × 10⁻¹³ mole m⁻² s⁻¹), respectively. Estimating the biomass with Equation (6) and with the lab rate at 6.17 × 10⁻¹¹ mole g⁻¹ s⁻¹ (see Seager et al. 2013), the biomass surface density would need to be about 3000 g m⁻², 80 g m⁻², and 0.001 g m⁻² for each stellar type, respectively. A globally averaged density of 3000 g m⁻² is likely too high; one of 80 g m⁻² is high but not impossible, according to terrestrial biodensities (see Seager et al. 2013).

The results show that CH₃Cl is a more viable biosignature gas in low-UV environments compared with high-UV environments. We emphasize that although our estimates of biomass surface density for Type III biosignature production are approximate, the resulting trend is robust.

For a spectral detection in transmission for our fiducial Earth transiting an M5V star, the required concentration is about 10 ppm in the wavelength range 3.2–3.4 μm. The surface bioflux and biomass estimates for a weakly active and quiet star, respectively, are 3.2 × 10¹⁵ molecule m⁻² s⁻¹ and 900 g m⁻² and 5.2 × 10¹¹ molecule m⁻² s⁻¹ and 0.001 g m⁻². The required biomass surface density for the weakly active M star is higher than the average surface biomass in Earth's oceans and the biomass surface density for the quiet M5V dwarf star is much lower than that of the Earth and is very plausible, again emphasizing the trend that low-UV radiation environments are more favorable for Type III biosignature gas accumulation.

The different values for the biosignature gas surface flux and the biomass estimates for transmission spectra compared with thermal emission spectra are in general due to longer atmospheric pathlengths and/or different favorable wavelengths (depending on the molecule of interest).

The fate of CH₃Cl in its destruction by H is to end up in its fully hydrogenated form, HCl, with the overall reaction as given by

$$\begin{equation} {\rm CH_3Cl + H_2 \rightarrow CH_4 + HCl}. \end{equation} \tag{ 25 }$$

HCl could be a bioindicator. The HCl molecule is stable against further photochemistry because if it is photolyzed, the Cl atoms generated will be predominately react with H to re-form HCl. HCl would not be expected to be present in significant levels at atmospheric altitudes for spectral detection without life taking non-volatile forms of Cl and putting them into the atmosphere, because all geological sources are non-volatile chlorides (such as NaCl) and any HCl that is volcanically released would be efficiently rained out of the troposphere. The limiting problem is that the HCl spectral features are too weak to be detectable and are likely to be contaminated by CH₄ in the 3–4 μm range.

As a second Type III biosignature gas example, we consider DMS. DMS has been studied before in oxidizing atmospheres by Domagal-Goldman et al. (2011), who concluded that DMS itself is not a potentially detectable biosignature gas in oxidized atmospheres under Sun-like UV radiation conditions, but one of its photolytic breakdown products, ethane, is detectable (we call this a "bioindicator" gas). Using the same atmosphere and framework as the above CH₃Cl example, we find a mixing ratio required for detection of 0.1 ppm in the 2.2–2.8 μm band (see Figure 5) for thermal emission spectra. Via photochemistry, this mixing ratio corresponds to a surface flux in our fiducial H₂-dominated atmosphere for a Sun-like star, a weakly active M5V dwarf star, and a quiet M5V dwarf star of 4.2 × 10¹⁹ molecules m⁻² s⁻¹ (6.9 × 10⁻⁵ moles m⁻² s⁻¹), 1.8 × 10¹⁹ molecules m⁻² s⁻¹ (3.0 × 10⁻⁵ moles m⁻² s⁻¹), and 2.4 × 10¹³ molecules m⁻² s⁻¹ (4.1 × 10⁻¹¹ moles m⁻² s⁻¹), respectively. Using a DMS lab production rate of 3.64 × 10⁻⁷ moles g⁻¹ s⁻¹ (Seager et al. 2013), we come up with an implied biomass surface density estimate for the three star types of about 200 g m⁻², 100 g m⁻², and 10⁻⁴ g m⁻², respectively. The first two values are high, but physically plausible compared with Earth's biomass surface density ranges. For transmission spectra, the numbers are about a factor of two higher for the weakly active and quiet M5V dwarf star (see Figure 7 and Table 3).

The DMS results show again that the lowest UV environment is the most favorable. There are two other relevant points related to DMS appearing to be a favorable biosignature gas in each of the three UV radiation environments studied. The first point is that gases destroyed by reactions with O (as opposed to gases destroyed by reactions with H) show a similar surface flux requirement between the Sun-like and weakly active M dwarf star. This is because the release of O from CO₂ photolysis is largely driven by Lyman α emission, which is similar at the habitable zones for the Sun-like and weakly active M dwarf stars used in this study (Figure 1).

The second point is that the high R_lab values used for DMS, and hence the low biomass surface density estimates, are a result of the biology of DMS production. On Earth, DMS is the waste product of consumption of Dimethylsulfoniopropionate (DMSP) by marine organisms consuming marine plankton. DMSP is accumulated in large amounts by some marine species. Thus, organisms that generate DMS do not have to invest their own resources to make DMS and so are not limited to how much they can make. Maximal production rates are therefore very high. This is discussed further in Seager et al. (2013).

In terms of a bioindicator, DMS will react with H₂ to generate CH₄ and H₂S. Neither is a useful bioindicator as CH₄ and H₂S are expected to be present in the atmosphere naturally. This is in contrast with oxidized atmospheres, where ethane may be a bioindicator gas, as expected for the end product of DMS photodestruction by the combination of methyl radicals generated from the attack of O on DMS (Domagal-Goldman et al. 2011).

As a third and fourth Type III biosignature gas example, we used CS₂ and OCS. For these two gases, we find the same trend as the other Type III biosignature gases: in a low-UV environment, the biosignature gases can accumulate (see Tables 2 and 3). The biomass estimates (as a plausibility check) are too high for the Sun-like and weakly active M dwarf stellar environments to be plausible compared with terrestrial biomass surface density values. OCS in the UV environment of a weakly active M dwarf star may be an exception with an estimate at the upper limit of plausibility.

As a fifth example, we describe N₂O. On Earth, N₂O is a Type I biosignature gas produced by nitrifying bacteria. N₂O is not likely to be produced in a thin H₂-rich atmosphere because there is unlikely to be much nitrate available. Here, we explain further. N₂O has been suggested as a biosignature gas in Earths atmosphere (Segura et al. 2005). N₂O is a Type I biosignature gas formed by two processes on Earth—the oxidation of ammonia by atmospheric oxygen and the reduction of nitrate in anoxic environments:

$$\begin{eqnarray} &&{\rm 2NH_3 + 2O_2 \rightarrow N_2O + 3H_2O}, \end{eqnarray} \tag{ 26 }$$

$$\begin{eqnarray} &&{\rm NO_3^- + H \rightarrow N_2O + H_2O}. \end{eqnarray} \tag{ 27 }$$

Analogous reactions on a hydrogen-dominated world would be the reduction of nitrate by atmospheric hydrogen:

$$\begin{equation} {\rm NO_3^- + H_2 \rightarrow N_2O + H_2O + OH^-} \end{equation} \tag{ 28 }$$

or the oxidation of ammonia by a geologically derived oxidant.

Nitrate is formed on Earth by the oxidation of NO generated by lightning in Earth's oxygen-rich atmosphere or by biological processes—neither are likely in an H₂-rich environment, so it is not clear whether nitrate reduction is a useful energy source in a world with an atmosphere rich in H₂. Ammonia oxidation requires a strong oxidizing agent, which again is likely to be missing from the environment.

N₂O as a Type I biosignature gas therefore seems unlikely, although not impossible, from very rare environments in which there are oxidized nitrogen species generated geochemically.

N₂O, however, could be a Type III biosignature gas as NO is for some organisms on Earth. We have calculated the surface fluxes for a detectable amount of N₂O in a thin, H₂-dominated atmosphere and find relatively low required surface fluxes. The reason is that N₂O is destroyed by photodissociation at a slower rate than by reaction with H. N₂O may therefore be a plausible biosignature gas candidate, even in an atmosphere subject to strong UV radiation (see Figures 6 and 7 and Tables 2 and 3). A biomass estimate (as a plausibility check) is not possible for N₂O, as it is only known as a Type I biosignature on Earth (and so therefore Type III R_lab rates are not available).

We have calculated biosignature gas accumulation in an atmosphere with 90% H₂ and 10% N₂ by volume. A super-Earth exoplanet atmosphere can have many other gas species. The concentration of the major destructive species (H, O, and OH) will depend on the amounts of these other gas species.

As an example, we explore the changing effect of the reactive species in an H₂-dominated atmosphere for different UV flux levels, based on the surface flux levels of CO₂ (Figure 2). A few key points are as follows. The H abundance is almost not affected by the CO₂ mixing ratios ranging from 10⁻⁸ to 10⁻². The O abundance depends on both CO₂ and UV levels, such that both a high CO₂ level and a high-UV flux lead to high atmospheric O. Only in extreme cases (e.g., H₂-dominated atmospheres with >1% CO₂, shown by dashed lines) is the abundance of O very close to the abundance of H. The OH abundance depends on a complex source-sink network, ultimately driven by H₂O and CO₂ photolysis. Notably, the amount of H is always at least four orders of magnitude larger than the amount of OH.

The effect of changing the H₂ mixing ratio and the addition and variation of other active gases on the H concentration will need to be considered in a case-by-case basis as they will react not only with H and OH but also with other gas species.

Whether or not a super-Earth planet can retain H₂ stably from atmospheric escape is not known. Although many models and studies for exoplanet atmospheric escape exist (see, e.g., Lammer et al. 2012 and references therein), the permanent limitation is that there are too many unknowns to provide a definitive and quantitative statement on which planets will retain H₂. One of the challenges is the unknown history and present state of the host star's EUV flux. Another major challenge is defining the mechanism for atmospheric escape for a given exoplanet, for example whether or not the regime of rapid hydrodynamic escape was reached in a planet's history or which non-thermal mechanism, if any was dominant (see Table 4.1, and references therein in Seager 2010). With an unknown initial atmospheric reservoir and an unknown present atmospheric composition, the regime and type of atmospheric escape is difficult to impossible to identify.

Some super-Earths will have been formed with atmospheres with H₂, based on both theoretical and observational evidence. Theoretically, planetary building blocks contain water-rich minerals that can release H (Elkins-Tanton & Seager 2008; Schaefer & Fegley 2010). Observationally, the large number and variety in radii of Kepler mini-Neptunes imply that these objects have an H or an H/He envelope to explain their radii. So, either from outgassing or the nebular capture of gases, some super-Earths should have started out with H₂-rich atmospheres and those with high enough gravity and low enough temperatures and/or magnetic fields should be able to retain the H₂ (e.g., Pierrehumbert & Gaidos 2011). Observational detections of H₂-rich atmospheres will ultimately be needed to confirm the scenario of thin H₂-dominated atmospheres on super-Earths.

Super-Earths with H₂-dominated atmospheres can have surface temperatures hotter than Earth due to an H₂ greenhouse effect from H₂–H₂ collision-induced opacities (Borysow 2002; Pierrehumbert & Gaidos 2011). While the hypothetical planets we have described in this paper were constructed to have 1 bar atmospheres with Earth-type surface temperatures, many H₂-dominated planet atmospheres are likely to have hotter surface temperatures than Earth, even for planets orbiting beyond 1 AU from their host stars.

An important question for understanding the potential of biosignature gases on a planet with an H₂-dominated atmospheres is therefore "how hot can a planet be and still sustain life?" On Earth, organisms that grow at 395 K are known (Lovley & Kashefi 2003; Takai et al. 2008) and have been cultured in the lab at elevated pressures equal to in situ pressures. Furthermore, proteins can function at 410–420 K (Tanaka et al. 2006; Sawano et al. 2007; Unsworth et al. 2007), motivating a consensus that life at 420 K is plausible (Deming & Baross 1993; Cowan 2004).

Life might exist at temperatures even higher than 420 K. The main argument for a maximum temperature for life involves the temperature at which the basic building blocks of life (DNA, proteins, carbohydrates, and lipids) break down. Many of the component chemicals of life, including DNA, many of the amino acids that make up proteins, and many of the key metabolites that allow lifes biochemistry to function, are rapidly chemically broken down above 470 K (e.g., Cowan 2004) The maximum temperature at which life could exist therefore may lie between 420 K and 470 K.

Many super-Earth atmospheres will be much more massive than the 1 bar atmosphere on Earth. For temperatures suitable for the existence of liquid water (see Section 5.3), the surface pressure could be as high as 1000 bar or higher (Wagner & Pruß 2002). There are three key points to show that the high surface pressures do not destroy the biosignature gases before they can reach the high atmosphere.

"Can life generate potentially detectable biosignature gases under a massive atmosphere?" The answer is yes, provided that the surface temperature is compatible with life. Then, in principle, life can survive and generate biosignature gases. The chemistry described in this paper still holds under a massive atmosphere, because the photochemical destruction occurs above 1 mbar. Furthermore, we showed in Seager et al. (2012) that the biomass surface density estimates are unchanged under a massive atmosphere as long as the photochemical loss rate dominates. For biosignature gases whose loss is dominated by deposition at the surface (i.e., are absorbed by the surface), then the biosignature source flux and hence biomass surface density will scale linearly with the planetary atmosphere mass.

The second key question is "can the high density and pressures on the surface under a massive atmosphere generate false positives?" The answer is almost entirely no, because we have shown that it is largely the Type III biosignature gases that are viable biosignature candidates. Recall that Type III gases are those produced for reasons other than energy extraction, and not from chemicals in the environment. Therefore, they are unlikely to be the product of nonbiological chemistry, a statement that holds even under a massive atmosphere. A comment related to NH₃ potential false positive is in order. For NH₃ to be generated from N₂ and H₂ kinetically, the temperature has to be well above any temperature compatible with life for any pressure where water is liquid, extrapolating from the known fact that at 300 bar and 673 K N₂ and H₂ still need a catalyst to be converted to NH₃ and higher pressures should not change this. The false positive risk is instead in detecting NH₃ without being able to identify the surface as cold enough not to possibly generate NH₃ kinetically.

The third key question is "will the high surface pressures enable fast chemical reactions that destroy the biosignature gases generated at the surface?" The answer is no, for pressures under about 1000 bar. In principle, if the upward diffusion or convective motions bring the biosignature gas to higher altitudes faster than the gas is destroyed by kinetic reactions, the higher surface pressures will not interfere with biosignature gas accumulation in the atmosphere.

Up to 1000 bar, reaction rates extrapolated from low-pressure kinetic experiments should be valid to an order of magnitude. An additional caveat is that low-pressure gas kinetics are usually measured at high temperature and we are extrapolating to high pressure and low temperature. For example, at 1000 bar and 300 K, we estimate the half life of hydrogenation of CH₃Cl to be 6.0 × 10¹¹ yr, the half life of DMS to be 3 × 10¹⁰ yr, and the halflife of N₂O to be 1 × 10²⁰ yr. These numbers are based on thermochemical equilibrium of H₂ and H based on the relative Gibbs free energy of formation of atomic hydrogen and T and P (e.g., Borgnakke & Sonntag 2009). The overview is that there is very little free H at high pressures and low temperatures since, in the absence of UV, H atoms will be generated almost entirely thermochemically.

At pressures above 1000 bar, we are less confident that chemistry can be extrapolated even qualitatively from low-pressure experiments. By 1000 bar, most gases will have densities approaching those of their liquids. Increases in pressure will force molecules closer than their van der Waals radii, directly altering molecular orbitals and reaction pathways. Below ∼1000 bar, we can consider molecules to be separate entities and we can still consider the molecules chemistry to be qualitatively similar to that of their dilute (ideal) gas state. Hence, order-of-magnitude extrapolations from low-pressure gas chemistry are justifiable.

Given the argument that life can generate biosignature gases on a planet with an H₂-rich atmosphere, but that the surface must have the right temperatures, how can we identify suitable planets for further study? The problem is that super-Earths are observed with a wide range of masses and sizes and we can anticipate that a range of atmosphere masses will also exist. A challenge is presented in observational atmosphere studies because we can only "see" to an optical depth of a few and this limiting optical depth can be reached well above any surface for a thick atmosphere. In many cases, surface conditions cannot be probed. Ideally, high-resolution spectra can be used to tell whether or not the atmosphere is thick or thin (i.e., whether or not one can observe down to the planetary surface), based on the shape of the spectral features, as described in detail in Benneke & Seager (2012, 2013).

We support the search for biosignature gases regardless of being able to classify a planet as habitable, because identifying biosignature gas molecules may be more readily attainable than high-spectral resolution characterization of a super-Earth atmospheric spectrum. That said, where possible, planetary radii can be used to discriminate planets worthy of follow up since those with small enough radii can be inferred as likely having thin atmospheres and those with radii large enough to have massive H₂ or H₂/He envelopes are unsuitable (Adams et al. 2008).

Biosignature gases can more easily accumulate in a low-UV radiation environment as compared with a high-UV radiation environment because the UV creates the destructive atmospheric species. We have shown this for H₂ atmospheres in this paper and Segura et al. (2005) have shown this for Earth-like planet atmospheres.

Whether or not truly UV-quiet M dwarfs exist and if UV activity is correlated with photometric stability is unclear. Recently, France et al. (2013) observed a small sample of six planet-hosting M dwarf stars with HST observations at far-UV and near-UV wavelengths and found none to be UV quiet. Other studies with much larger numbers of M dwarf stars are ongoing, including some with UV emission from the Galaxy Evolution Explorer (A. West, 2013, private communication). A general understanding is that magnetic activity, as traced by Hα in M dwarfs, decreases with age but that M dwarfs appear to have finite activity lifetimes such that the early-type M dwarfs (M0–M3) spin down quickly with an activity lifetime of about 1–2 Gyr whereas later type M dwarf stars (M5–M7) continue to spin rapidly for billions of years (West et al. 2006, 2008).

For the time being, UV (that is, the relevant FUV and EUV) radiation emitted by the stars of interest is not usually measured or theoretically known and so we have worked with three different UV radiation environments (Figure 1).

A relevant point for quiet M stars with extremely low-UV radiation (if they exist) is that false positives for biosignature gases destroyed by photolysis may also more easily accumulate. This is relevant for NH₃, for which the lifetime in our fiducial H₂ planet atmosphere for a planet orbiting a quiet M star is about 1.4 Gyr, according to our photochemistry models. This means that the false positive risk that comes from primordial NH₃ would be high.⁹ Related issues with other gases that are primarily destroyed by photolysis (and not destroyed by reactions with H and O) should be investigated.

Is there any hope that the next space telescope, JWST could be the first to provide evidence of biosignature gases? Yes, if–and only if–every single factor is in our favor.

First, we need to discover a pool of transiting planets orbiting nearby (i.e., bright) M dwarf stars. Second, the planet atmosphere should preferably have an atmosphere rich in molecular hydrogen to increase the planetary atmosphere scale height. Third, the M dwarf star needs to be a UV-quiet M dwarf star with little EUV radiation. Fourth, the planet must have life that produces biosignature gases that are spectroscopically active.

Several biosignature gases, if they exist, are detectable with tens of hours of JWST time, based on our detection metric. Although our detection metric assumes that photon noise is the limiting factor, many more detailed simulations of JWST detectability show that spectral features of a similar magnitude are detectable (e.g., Deming et al. 2009).

For detecting molecules using transmission spectroscopy, the background exoplanet atmosphere dominated by H₂ or CO₂ has little effect on the detectability of the biosignature gases of interest that we studied. This is because the transmission observations are better performed in the near-IR than in the mid-IR because of a higher stellar photon flux at near-IR wavelengths and the contamination effects of either the dominant CO₂ absorption or collision-induced H₂–H₂ absorption are minimal in the near-IR. As long as all of the biosignature gases of interest have features in the near-IR (see Figure 7 for the spectral features), these gases may be detected for atmospheres with any level of CO₂. The key issue here, instead of spectral contamination, is the mean molecular mass. The depth of the transmission spectral feature is one order of magnitude larger for H₂-dominated atmospheres compared with CO₂-dominated atmospheres (see the scale height in Equation (1)).

For detecting molecules via thermal emission with future direct imaging techniques, one may expect the CO₂ or H₂–H₂ contamination to be important because the thermal emission of a planet peaks in the mid-IR where CO₂ and H₂–H₂ contamination is most substantial. For individual gases, however, there are often multiple absorption bands to mitigate this issue. Similarly, a variety of wavelength ranges are usually available to choose from for the other biosignature gases of interest studied in this paper (as shown in Figures 4–7).

At this point, we conclude by emphasizing a related point that the plausibility of a specific biosignature gas depends on the planet surface gravity, atmospheric pressure, and other characteristics, because such characteristics affect which atmospheric wavelength "windows" are most favorable. Individual planets and their atmospheres should be considered on a case-by-case basis.