Introduction

Present-day Mars is extremely cold and dry, but many geomorphological and geochemical evidence, such as valley networks, suggest an active water cycle at 3.8–3.6 Ga1. The detection of phyllosilicates over the Noachian terrain also supports the existence of widespread liquid water on early Mars2. The habitability of Mars has been of great interest that triggered previous and ongoing Martian explorations. Although water is a probable requirement of Martian habitability, this molecule is not an organic compound in genomic and catalytic bio-molecules that supports the fundamentals of life. Investigations on the organic synthesis on ancient Mars fill the gap by verifying the possibility, environment, and age of the chemical evolution to potential ancient Martian life.

Formaldehyde (H2CO) is simple organic matter that can be formed through various chemical reactions in planetary atmospheres. H2CO is a highly soluble and reactive molecule, and thus has the potential to play a significant role in the abiotic formation of bioorganic molecules3. For example, amino acids are formed by reactions involving H2CO, NH3, and HCN via the Strecker reaction4. Additionally, in ammonia-involving formose-type reactions, the condensation of H2CO with NH3 yields various amino acids5,6. The formose reaction is a thermally driven aqueous process that generates many sugars from H2CO, including ribose, a fundamental building block of RNA that is regarded as a key molecule for the origin of life7,8,9. Therefore, determining whether the surface environments on early terrestrial planets fostered the production of H2CO is crucial for understanding prebiotic chemical evolution to the origin of life.

The evidence that supports the existence of liquid water have led many scientists to imagine a warm early Martian climate as it is on Earth1,2. Previous numerical studies have attempted to reproduce warm early Mars; however, 3-D global circulation model (GCM) studies have not been able to reproduce the continuous presence of liquid water on the surface with CO2-H2O atmospheres10,11. To reconcile this geomorphological evidence, episodic melting scenarios driven by the supply of reducing gases through volcanic outgassing or meteorite impacts have been proposed12,13,14,15.

Pinto et al.16 provided an estimate for the photochemical production of H2CO in the N2-dominated atmosphere of primitive Earth, with the number density of H2CO near the surface approximated to be ~ 108 cm−3. A similar amount of H2CO production was also predicted for early Earth conditions by Harman et al.17. However, the production of H2CO in a CO2-dominated atmosphere on early Mars has not yet been thoroughly investigated. Understanding its production on early Mars can provide insights into the potential of life on the planet.

In this study, we investigated the production of H2CO in a thick CO2-dominated atmosphere containing H2 and CO on early Mars using a 1-D photochemical model. We then estimated ribose production using our simulation results and experimental data.

Methods

Model description

To calculate the atmospheric production of H2CO, we adapt a one-dimensional photochemical model, PROTEUS (Photochemical and RaiatiOn Transport model for Extensive USe), detailed by Nakamura et al.18, for early Martian conditions. This model has been successfully applied to other planetary atmospheres, such as the Jovian ionosphere19 and present-day Martian atmosphere20. It solves the continuity equations involving chemical reactions and vertical transport until the profiles of each species reach a steady state. We consider 63 chemical reactions (Supplementary Table S1 online) for 8 neutrals: CO2, CO, H2, H2O, O2, H2O2, O3, H2CO, 6 radicals: H, O, OH, HO2, O(1D), HCO, and an ion of CO2+ in a 2-bar CO2-dominated atmosphere. Though the atmospheric surface pressure of early Mars is still not well constrained, 3-D global circulation model studies suggest that a 2 bar CO2 atmosphere with a few percentages of reducing gas is required for a warm climate14,15. The exobase altitude is defined as a pressure level of 10−9 mbar. We utilize the H2O vapor number density and temperature profiles up to ~ 60 km (Fig. 1) from the 2-bar global mean results with an obliquity of 40° computed using a 3-D paleo-Mars global climate model15,21. The global climate model assumed three scenarios of an atmosphere containing 0, 3, and 6% H2. For the H2O density profiles above ~ 60 km, we assume the same mixing ratio up to the exobase, assuming the effect of cold trap. The sensitivity to the H2O vapor content in the atmosphere is discussed in Results section. For temperature, we assume isothermal up to the lower boundary of the thermosphere (6 × 10−3 Pa) considering radiative equilibrium and then extrapolate it into the thermosphere22,23. The exobase temperature is fixed at 800 K, corresponding to 10 × EUV at 3.8 Ga24. We adopt the solar spectrum from 3.8 Ga estimated by Claire et al.25. We use the updated H2O absorption cross section measured by Ranjan et al26. We assume the up-to-date absorption cross sections for all the species to the best of our knowledge. References for cross sections of all species at each wavelength are presented in Nakamura et al.18. The CO2+ concentration profile calculated by the ionosphere photochemistry model27 is fixed at the same pressure altitude as present-day Mars to represent the dissociation reaction of H2 with CO2+ to produce atomic H in the upper atmosphere, allowing us to calculate the escape flux of hydrogen23. Although the density of CO2+ in the early Martian atmosphere has uncertainty, its profile has little impact on the result, because the escape flux of hydrogen is limited by H2 diffusion from the lower atmosphere. We adopt eddy diffusion coefficient profiles using the typical formula adapted for other planets than Earth28. The vertical temperature, H2O number density, and eddy diffusion coefficient profiles are shown in Fig. 1. The temperature and water vapor profiles of the H2 0% scenario shown in Fig. 1 are used to calculate the H2CO production under the conditions of 0.1, 0.01, 0.001, and 0.0001% H2.

Figure 1
figure 1

Background atmospheric conditions in early Mars. Temperature-altitude profile (a and c), Eddy diffusion coefficient profile (b), and H2O number density profile (d). The solid, dashed, and dash-dotted lines correspond to H2 6%, 3%, and 0% conditions, respectively. The temperature and H2O number density profiles of the H2 0% case in this figure are used to calculate the results for H2 0.1, 0.01, 0.001, and 0.0001% cases.

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For the upper boundary, we assume Jeans escape of H and H2 and fix the O escape rate at 2.6 × 108 cm−2 s−1 corresponding to 10 × EUV conditions24. The O escape rate does not have a large impact on the results because it does not control O2 abundance in the lower atmosphere where H2CO is dominantly produced. Deposition velocities are applied to H2O2, HO2, H2CO, HCO, OH, O, and H29,30. The deposition velocities of each species are shown in Supplementary Table S2 online. We compute the dry deposition of H2CO by imposing the deposition velocity. This model also includes the rainout of H2CO throughout the atmosphere using the same parameterization as Hu et al28:

$${\text{k}}_{{\text{R}}} \left( {\text{z}} \right) = {\text{f}}_{{\text{R}}} \times \frac{{{\text{n}}_{{{\text{H}}_{2}{\text{O}}}} \left( {\text{z}} \right){\text{k}}_{{{\text{H}}_{2}{\text{O}}}} \left( {\text{z}} \right)}}{{55{\text{N}}_{{\text{A}}} \left[ {{\text{L}} \times 10^{ - 9} + \left( {{\text{H}}^{\prime } {\text{RT}}\left( {\text{z}} \right)} \right)^{ - 1} } \right]}},$$
(1)

where kR is rainout frequency, fR is a reduction factor which is an adjustable parameter to represent the reduction relative to Earth’s hydrological cycle, \({\text{n}}_{{\text{H}}_{2}{\text{O}}}\) is the number density of H2O, \({\text{k}}_{{\text{H}}_{2}{\text{O}}}\) is the precipitation rate assumed to be 2 × 10–6 s−1, NA is Avogadro’s constant, L is the liquid water content assumed to be 1 g m−3, H′ is the effective Henry’s Law constant assumed to be 1.3 × 104 M atm−1 taken from Giorgi & Chameides31, R is gas constant, and T is temperature. The rainout rate is then obtained by multiplying kR by number density of H2CO. We assume fR to be 1, assuming that early Mars had a hydrological cycle similar to Earth’s. Sensitivity to fR is discussed in Results section. The boundary conditions for H2 and CO are fixed in the calculation of H2CO production in the results. We impose H2 outgassing and CO deposition velocity as a boundary condition to determine the possible range of H2 and CO mixing ratio in the following section.

Background H2 and CO atmospheric conditions

Potential sources of H2 gas on early Mars include volcanic degassing29, meteorite impacts32,33,34, and serpentinization35. The upper limit of the H2 outgassing rate is estimated to be 8 × 1011 cm−2 s−1 considering the supplies from volcanism and serpentinization29. We computed H2 mixing ratios in a background 2-bar CO2 atmosphere for various H2 outgassing rates to investigate the plausibility of CO2 atmospheres enriched in H2, assuming a fixed CO deposition velocity of 10−8 cm s−129 (see Supplementary Fig. S1 online). An H2 outgassing rate of ~ 5 × 1011 cm−2 s−1 yields a 5–6% H2 mixing ratio in a 2-bar CO2 atmosphere. This result is consistent with that of Batalha et al.29, in which an H2 outgassing rate of 8 × 1011 cm−2 s−1 yields a ~ 5% H2 mixing ratio in a 3-bar CO2 atmosphere. The minimum value of the H2 mixing ratio is ~ 1 × 10–6 when assuming the absence of H2 degassing. Furthermore, Chassefiere et al.35 suggested that serpentinization-derived CH4 trapped in the cryosphere could have been released into the atmosphere, producing a transient 1–2 bar CO2 atmosphere containing 10–20% H2 gas. Thus, a 6% H2 mixing ratio in a background 2-bar CO2 atmosphere is plausible. This study considers a range of H2 mixing ratios from 1 × 10−6 to 0.06.

The abundance of CO on early Mars is not well constrained. A dense and cold CO2 atmosphere is likely to enter the CO runaway state23. The CO and O liberated from CO2 photolysis no longer recombine because of the lack of odd hydrogen species that catalyze CO2 recombination. Moreover, because the equilibrium timescale of CO is relatively long, ranging from several million to several hundred million years23, it is insufficient to consider only a single steady state condition. Therefore, it is necessary to vary the CO mixing ratio as a parameter within a specific range.

To determine the possible CO range, we calculated the CO mixing ratios over a wide range of CO deposition velocities for 0, 3, and 6% H2 cases. Different temperatures and H2O profiles are used for each H2 case obtained from the GCM results15. We assume a free lower boundary condition for H2 in the 0% H2 case, whereas the number density is fixed for the 3 and 6% H2 conditions.

As a result, the 0% H2 case enters a CO runaway state with a low CO deposition velocity of < 10−10 cm s−1 (see Supplementary Fig. S2 online) as suggested by the previous photochemical model study of early Mars conditions36. Conversely, warmer atmospheres containing 3 or 6% H2 have a maximum CO mixing ratio of ~ 1% because sufficient H2O vapor promotes CO2 recombination. The CO deposition velocity on an abiotic ocean planet is estimated to be 10−9–10−8 cm s−137,38; however, there is no known efficient process to remove CO at the surface without the ocean28. Considering that ancient Mars experienced episodic cold and warm climates1, a CO2 atmosphere would have been in a CO runaway state during the ice age36, while the atmosphere would have been stable with a minimum CO mixing ratio of 1% in warmer climates. Based on these results, the possible range of CO on early Mars should be from 1 to 50%.

Results

Photochemical production of formaldehyde

The present 1-D photochemical model shows that H2CO forms at number densities of ~ 5 × 109 and ~ 7 × 10−2 cm−3 near the surface under 6% H2 condition and 0.01% H2 condition, respectively (Fig. 2). Production in a 6% H2 atmosphere is substantial; approximately 50 times higher than that estimated for Earth's Hadean atmosphere16.

Figure 2
figure 2

Number density profiles of main species under the 6% H2 (solid line) and 0.01% H2 (dashed line) conditions. CO is fixed at 1% in both cases.

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H2CO is predominantly formed through a radical–radical reaction of two HCO molecules39:

$${\text{HCO}} + {\text{HCO}} \to {\text{H}}_{2} {\text{CO}} + {\text{CO}}{.}$$
(R1)

The dominant production path of HCO is a three-body reaction of H and CO:

$${\text{H}} + {\text{CO}} + {\text{M}} \to {\text{HCO}} + {\text{M,}}$$
(R2)

where M represents the background gas, mainly CO2 in this model. In the early Martian atmosphere, H and CO in R2 are derived from H2O and CO2 photolysis, respectively. HCO is dominantly destroyed by a reaction with O2:

$${\text{HCO}} + {\text{O}}_{2} \to {\text{HO}}_{2} + {\text{CO}}{.}$$
(R3)

These reactions imply that the production of H2CO decreases with O2 because O2 reacts with HCO through R3, thereby reducing the rate of R1.

The H2CO production in the 6% H2 atmosphere is notably higher than that under lower H2 conditions. This results from the discrepancy in the O2 abundance in the lower atmosphere between these two conditions. There is a sharp decline in the O2 density below 60 km for 6% H2 (Fig. 2). The lack of O2 in the 6% H2 atmosphere decreased the reaction rate of HCO with O2 (R3). Consequently, a larger amount of HCO remains near the surface, ultimately increasing the amount of H2CO via reaction R1. Two mechanisms below contribute to the decrease in O2.

The first mechanism is driven by H atoms at altitudes greater than 100 km. The source of O2 near the surface is the downward transport of O2 liberated from CO2 photolysis at high altitudes. The significant difference between the 6% and 0.01% H2 cases is the number of H atoms at altitudes above 100 km. H2 is transported upward and photolyzed into H atoms by the solar UV flux, producing more H in the 6% H2 case. O2 reacts with H at high altitudes and is converted back into CO2 via the following reaction:

$${\text{H}} + {\text{O}}_{2} + {\text{M}} \to {\text{HO}}_{2} + {\text{M}}$$
(R4)
$${\text{H}} + {\text{HO}}_{2} \to 2{\text{OH}}$$
(R5)
$${\text{CO}} + {\text{OH}} \to {\text{CO}}_{2} + {\text{H}}{.}$$
(R6)

The altitude profiles of the reaction rate of R4 between the 6% and 0.01% H2 cases indicate that R4 shifted upward under the 6% H2 condition (Fig. 3). The loss of O2 by H atoms through R4 above ~ 150 km results in a decrease in O2 downward flux below ~ 135 km, as shown in Fig. 4.

Figure 3
figure 3

A comparison of O2 loss reaction rates per one O2 molecule (= reaction rate / O2 number density) under the 6% H2 (solid line) and 0.01% H2 (dashed line) conditions. The black and blue lines show the reaction rates of \({\text{H}} + {\text{O}}_{2} + {\text{M}} \to {\text{HO}}_{2} + {\text{M}}\) (R4) and \({\text{HCO}} + {\text{O}}_{2} \to {\text{HO}}_{2} + {\text{CO}}\) (R3), respectively.

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Figure 4
figure 4

A comparison of O2 downward flux profiles under the 6% H2 (solid line) and 0.01% H2 (dashed line) conditions.

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The second mechanism is driven by the HCO catalytic cycle. In the present-day Martian atmosphere, odd hydrogen species act as catalysts to recombine CO and O into CO2 through the following cycle40:

$${\text{H}} + {\text{O}}_{2} + {\text{M}} \to {\text{HO}}_{2} + {\text{M}}$$
(R4)
$${\text{O}} + {\text{HO}}_{2} \to {\text{OH}} + {\text{O}}_{2}$$
(R7)
$${\text{CO}} + {\text{OH}} \to {\text{CO}}_{2} + {\text{H}}$$
(R6)

Net: \({\text{CO}} + {\text{O}} \to {\text{CO}}_{2}\).

The net result is the recombination of CO2.

In a dense CO2 atmosphere, HCO catalytic reactions are responsible for converting H to HO2 in addition to R4:

$${\text{HCO}} + {\text{O}}_{2} \to {\text{HO}}_{2} + {\text{CO}}$$
(R3)
$${\text{H}} + {\text{CO}} + {\text{M}} \to {\text{HCO}} + {\text{M}}$$
(R2)

Net: \({\text{H}} + {\text{O}}_{2} \to {\text{HO}}_{2}\).

This is because a larger amount of CO is formed in a denser atmosphere owing to the lack of amounts of odd hydrogen, while O2 is more abundant than CO in the present-day Martian atmosphere. The reaction rate profile of R3 in a 6% H2 atmosphere indicates that the HCO catalytic cycle dominates at ~ 50–60 km (Fig. 3), where the O2 density declines sharply (Fig. 2). This cycle significantly reduces O2 in the lower atmosphere below 50–60 km.

Deposition of formaldehyde

A 6% H2 mixing ratio enables the presence of an ocean in a warm climate15. In a warm environment with an ocean and a CO deposition velocity of 10−8–10−9 cm s−137,38, the CO mixing ratio is ~ 1%, as shown in the 3% and 6% H2 cases in Supplementary Fig. S2. Consequently, the deposition flux of H2CO into the surface liquid water is approximately 3 × 109 cm−2 s−1 in a warm climate.

When the mixing ratio of H2 and CO decreases, the deposition of H2CO decreases. However, the H2CO decrease is not gradual, and there is a respective large drop with the decrease in H2 and CO. Large amounts of H2CO are deposited in an atmosphere containing either > 0.1% H2 or > 50% CO (Fig. 5).

Figure 5
figure 5

H2CO deposition flux as a function of CO and H2 mixing ratios assuming a 2-bar CO2 background atmosphere.

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The high-altitude H generates a large drop between 0.01 and 0.1% H2, as shown in Fig. 5. The large decrease between 10 and 50% CO is associated with the HCO catalytic cycle (Fig. 5). Increasing the CO produces more HCO (R2), thereby removing O2 in the lower atmosphere (R3). In an atmosphere containing 10% CO, a steady state is reached with more O2 and less HCO. Conversely, the atmosphere containing 50% CO reaches a steady state with less O2 and more HCO. Therefore, in an atmosphere with H2 < 0.01%, more H2CO is produced with 50% CO.

Formaldehyde deposition in more abundant H2 and H2O conditions

We first assessed the impact of a higher H2 mixing ratio than 6% on H2CO production, with the CO mixing ratio fixed at 1% and H2O and temperature profiles from the 6% H2 GCM results15. This assumption may not represent early Martian conditions, but it allows the estimation of the upper limit of H2CO production in a CO2-dominated atmosphere on early Mars and provides insights into an exoplanet analog for a CO2-dominated atmosphere enriched with H2. As shown in Fig. 6, increasing the H2 mixing ratio from 1 to 20% increased the H2CO deposition slightly. However, the deposition flux does not exceed 5 × 109 cm−2 s−1. This result implies that H2CO production under early Martian atmospheric conditions would be close to its maximum in H2-rich CO2-dominated atmospheres.

Figure 6
figure 6

H2CO deposition fluxes as a function of the H2 mixing ratio (a) and H2O factor (b). The temperature profile of the 6% H2 case is used for both calculations. The bottom panel (b) shows H2CO deposition fluxes obtained by multiplying the H2O profile of the 6% H2 case by a factor ranging from 0.01 to 100.

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Subsequently, we multiplied the H2O profile of the 6% H2 case by a factor ranging from 0.01 to 100 while maintaining the H2 and CO mixing ratios at 6% and 1%, respectively, with the temperature profile of the 6% H2 case. This approach accounts for spatiotemporal variations in water vapor content on a global scale. H2O vapor changes by ~ 10 times over the global scale in the global climate model results21. Considering the seasonal difference, changes by ~ 2 orders of magnitude are reasonable for early Mars conditions. As an upper and lower limit, we changed the amount of H2O vapor by 4 orders of magnitude. This parameter survey may also be useful for exoplanets’ environment. A higher water vapor content significantly increases H2CO production (Fig. 6b). The H2CO deposition flux increases up to ~ 1 × 1012 cm−2 s−1 with a 100 times higher H2O density. Under conditions of abundant H2O, additional H atoms are generated through H2O photolysis, resulting in increased HCO formation through R2 and, consequently, greater H2CO production through R1. In addition, an increase in precipitation also contributes to the increased rainout rate of H2CO. This indicates that in a warm climate with 6% H2, the limiting factor for H2CO production is not H2 but the abundance of H2O. The timescale for the change in H2CO formation due to variations in water vapor content is a few years in our calculation. This result suggests that H2CO production on early Mars exhibited local variations depending on the availability of water vapor. However, further studies are needed to clarify this effect, as horizontal transport may mitigate differences in H2CO abundance.

We also investigated the effect of the reduction factor fR in Eq. (1) on H2CO deposition flux. The calculated H2CO deposition fluxes with a reduction factor of 0.1, 0.5, and 1 are shown in Supplementary Fig. S4. In this calculation, H2 and CO mixing ratios are fixed at 6% and 1%, respectively, with temperature and H2O profiles of the 6% H2 case. When fR is set to 0.1, it is reduced to 8 × 108 cm−2 s−1, approximately 1/4 of the value when fR is 1. Although it is not yet constrained, GCM results suggest that globally averaged precipitation on early Mars may have been 10 times smaller than on Earth15. Further modeling studies combining photochemistry with GCM would be helpful for a more accurate estimation of the rainout rate.

Discussion

Formation of formaldehyde throughout Mars’ history

The deposition rate of H2CO reaches the order of 108 or 109 cm−2 s−1 under conditions where the mixing ratio of H2 is higher than 0.1%, regardless of the CO mixing ratio (Fig. 5). The H2 mixing ratio of 0.1% is equivalent to ~ 1010 cm−2 s−1 of H2 outgassing rate in a steady state (Supplementary Fig. S1). This rate is comparable to the estimated H2 degassing rates on present-day Earth, which has an upper mantle of quartz-fayalite-magnetite (QFM) oxidation buffer29,41. The H2 degassing rate on early Mars is unclear. However, Martian meteorites suggest that the mantle was more reduced than the Earth’s upper mantle, with oxygen fugacity around the iron-wüstite (IW) buffer42,43. Given that the oxygen fugacity of the Martian upper mantle was buffered near IW + 1, the H2 degassing rate is estimated to be ~ 1011 cm−2 s−129,41. This suggests that the large atmospheric production of H2CO continued during past periods of active volcanic degassing, regardless of the CO mixing ratio (Fig. 7). An increase in the H2 mixing ratio from 0.1 to 6% increases the H2CO deposition rate by approximately 10 times the global mean value (Fig. 5). This increase was mainly due to the increase in the number density of H2O by 10 times due to the warming effect of increasing H2 (Fig. 1d). The number density of H2O in the atmosphere also differs (e.g., 10 times) depending on the local availability of H2O15. This difference could provide local H2CO deposition rates on Mars that are several tens of times higher and lower than the global mean value (Fig. 6b).

Figure 7
figure 7

Scenario for the atmospheric H2CO production at ca. 3.8–3.6 Ga (top panel), ca. 3.5–3.0 Ga (middle panel), and after ca. 3.0 Ga (bottom panel). In the Noachian and early Hesperian periods (3.8–3.6 Ga), the synthesized H2CO in the atmosphere was deposited into the ocean, forming bio-important molecules, such as ribose. In the middle and late Hesperian (3.6–3.0 Ga), H2CO was sporadically formed. Even in the period when H2CO was abundantly formed, subsequent formose reaction does not proceed due to the acidic condition of the water. From the Amazonian to the present (after 3.0 Ga), the production of H2CO is deficient, as in the case with H2 < 0.1% in the middle and late Hesperian.

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Volcanic degassing would have decreased from the late Hesperian to the Amazonian, decreasing the mixing ratio of H2 and H2O in the atmosphere42. A decrease in the H2 mixing ratio to below 0.1% dramatically decreased the H2CO deposition rate by a factor of 10−10 (Fig. 5). Thus, the late Hesperian to early Amazonian was a transitional period from a high to a meager H2CO deposition rate (Fig. 7).

Formation of organic compounds in the ocean

The continuous conversion of CO2 and CO into highly soluble H2CO in the early Martian atmosphere may have transferred carbon from the atmosphere to the ocean. Another mechanism that converts atmospheric carbon into H2CO involves iron-rich asteroids/meteorites44. Such impacts might have formed H2CO, both locally and temporally. The overall impact-induced production would have been smaller than the continuous global production of H2CO in the atmosphere through the photochemical reactions presented in this study (see Supplementary text online). Another source of H2CO discussed previously is the oxidation of CH4 with iron oxide, which was proposed to explain the tentative detection of H2CO on present-day Mars45.

H2CO is highly reactive. Carbon transferred from the atmosphere as H2CO could further be converted into various organic compounds. One of the most well-known reactions is the formose reaction, in which formaldehyde oligomerizes to form various sugar molecules in alkaline solutions7. A recent study found that this type of reaction can form sugars, including ribose, even in neutral solutions46. To estimate ribose production in the early Martian ocean, we make the following assumptions: one-third of the surface area was covered by an ocean as suggested by the analysis of the distribution of delta and valleys47, the seawater pH at ~ 3.8 Ga in early Mars was near-neutral, as the late Noachian marked a transition period from alkaline to acidic water pH48, early Mars had a hydrological cycle similar to Earth’s (fR = 1), the atmosphere contained the same fraction of glycolaldehyde as an estimated atmosphere of the early Earth17, and the conversion rate of H2CO into ribose was ~ 3.5 × 10−6 molRib molFA−1 estimated by the formose-reaction experiment46. By combining the calculated H2CO deposition flux of 3 × 109 cm−2 s−1 (= 1 × 1021 m−2 yr−1), the annual ribose production in the ocean Prib is estimated to be 4 × 104 kg yr−1 as follows:

$${\text{P}}_{{{\text{rib}}}} = \frac{{{\text{f}}_{{{\text{H}}_{2} {\text{CO}}}} {\text{S}}_{{{\text{sea}}}} {\text{Y}}_{{{\text{rib}}}} {\text{M}}_{{{\text{rib}}}} }}{{{\text{N}}_{{\text{A}}} }},$$
(2)

where \({\text{f}}_{{\text{H}}_{2}{\text{CO}}}\) is the deposition flux of H2CO (m−2 yr−1), Ssea is the area covered by the ocean (m2), Yrib is the conversion rate of H2CO into ribose (mol mol−1), Mrib is the molar mass of ribose (kg mol−1), and NA is Avogadro’s constant. When fR is 0.1, the annual ribose production is estimated to be 1 × 104 kg yr−1. These suggest that bio-important sugars including ribose might have been continuously formed in water bodies on the surface. In this assumption, we disregarded several H2CO consumption processes in seawater, including photolysis, hydrolysis, and reactions with other reactive molecules such as ammonia3,49. It is unclear whether ammonia was present on early Mars, but it may have transiently been present in the atmosphere or in water on early Mars, potentially due to processes, such as episodic volcanic degassing or impact degassing. Previous studies on early Earth indicated that reducing gases, including ammonia, could be generated through impact events50,51. This process may have analogously occurred on early Mars. In the presence of ammonia, the formose reaction forms various nitrogen containing organic matter including proteinogenic amino acids5. Nitrogen containing organic matter has been found in a Martian meteorite52. The formose reaction also forms refractory organic matter, similar to cometary and meteoritic insoluble organic matter6. The photochemical H2CO calculated in this study and its following formose reaction may be related to the origin of refractory and non-refractory organic matter found in the 3.5-billion-year-old lacustrine mudstones of Mars53 and Noachian carbonates in a Martian meteorite52. However, it is still difficult to distinguish whether this organic matter was derived from the photochemical H2CO. One way to distinguish them is to compare their carbon isotope compositions. The carbon isotopic analysis onboard the Curiosity rover detected an anomalously depleted 13C in organic matter54. The deposition of photochemical H2CO which experienced CO2 photolysis-driven carbon isotope fractionation might explain this depletion55. Our future work is to include carbon isotope fractionation in the model and compare it with the isotope observation data. Significant H2CO synthesis on the warm Noachian Mars with liquid surface water allowed for the formation of sugars and amino acids (Fig. 7). The formation of H2CO would have sporadically continued on Hesperian to early Amazonian Mars, but the transition to an ice-covered and acidic surface environment on Hesperian Mars dramatically decreased the possibility of the formation of the building blocks of life, because the production of sugars and amino acids through the formose-type reaction substantially decreased in acidic water46,48,56,57. Therefore, the time period suitable for the formation of bio-important molecules on early Mars might be limited to the Noachian and potentially early Hesperian Mars, the warm climate era before the pH of the surface liquid water became acidic.

Secondary concentration processes are essential for synthesizing bio-important molecules on planetary surfaces. Early Mars may have experienced episodic warm and cold climate periods33. During the transition from warm to cold periods, a large amount of oceanic water would have been stored as snow on land; thus, oceanic water would have become concentrated. Such evaporative environments might further promote chemical evolution to form biopolymers, such as proteins and RNAs because the primary reaction that forms these molecules is dehydration reactions58,59. These reactions may have been promoted by carbonate and borate, which have been shown to be present on Mars60,61,62,63,64. Future studies considering topography and secondary concentration processes are important to elucidate the possibility of RNA synthesis on early Mars.

Conclusions

Formaldehyde (H2CO) is a crucial organic matter in the formation of bioorganic molecules such as amino acids and ribose. We investigated the atmospheric production of H2CO on early Mars using a one-dimensional photochemical model. We assume a 2-bar background CO2-dominated atmosphere with various concentrations of H2 and CO while adopting temperature and H2O profiles from a 3-D paleo-Mars global climate model. Our results show that a larger amount of H2 leads to a more significant production of H2CO owing to the reduction in O2 abundance in the lower atmosphere. Two mechanisms cause this O2 reduction: (1) the chemical reaction with H atoms at high altitudes above 100 km, resulting in a decrease in O2 downward flux, and (2) the HCO catalytic cycle at ~ 50–60 km reducing O2 in the lower atmosphere below 60 km. In a warm climate, the number density of H2CO is ~ 5 × 109 cm−3 near the surface, and its deposition into the ocean is 3 × 109 cm−2 s−1 assuming that early Mars had a hydrological cycle similar to Earth’s. The sensitivity analysis of water vapor implies that H2CO production could have varied locally in correlation with the abundance of water vapor. Our results suggest that a continuous supply of H2CO could be used to form various organic compounds, including life's building blocks, such as amino acids and sugars. This photochemically produced H2CO could be a possible origin for the organic matter observed on the Martian surface. Given the previously reported conversion rate from H2CO to ribose, the calculated H2CO deposition flux suggests a continuous supply of bio-important sugars on Noachian and early Hesperian Mars.