4.2.1 Interannual Variability

The 1300–2200 ppmv abundances of oxygen measured by SAM are generally in the same range as prior measurements that span different sets of in situ and remote sensing observations of O 2 in the upper atmosphere and at the surface of Mars. The repeated measurements of SAM provide the most robust measurements to date, as the previously reported values have been made with substantial uncertainty and limited frequency, besides being in different regions of the atmosphere and the surface. From five sets of mass spectral measurements obtained within days of landing of Viking lander 1 (VL1) on 20 July 1976, Owen and Biemann (1976) reported an O 2 mixing ratio in the 1000–4000 ppmv range. Owen et al. (1977) subsequently concluded that the oxygen measurements had considerable scatter of a factor of 2, which they attributed to instrumental causes. It seems likely, as England and Hrubes (2004) suggest, that Owen et al. (1977) based the 1300 ppmv value for O 2 mixing ratio given in their Table 1 on measurements reported by Barker (1972) and Carleton and Traub (1972), which were obtained with high‐resolution 762 nm O 2 ‐band spectroscopy from terrestrial observatories (see also Trauger & Lunine, 1983 who report a slightly lower value from ground‐based spectroscopy: ~1200 ppm when scaled to 6 mbar surface pressure). The Viking GEx experiment reported an upper limit of 1500 ppmv (Oyama & Berdahl, 1977). England and Hrubes (2004) calculated a seasonal variation in O 2 between 2,500 and 3,300 ppmv, by inverse scaling of O 2 to atmospheric pressure measured by the Viking. It is important to note that the reference O 2 value they used in their scaling calculation (3,000 ppmv) is the value VL1 measured at 125–300 km, in the upper atmosphere above the homopause (Nier & McElroy, 1977), which is not representative of O 2 at the surface. Mars Express SPICAM observed 4,000 ppmv O 2 averaged over 90–130 km and six observations (Montmessin et al., 2017; Sandel et al., 2015). Any seasonal or temporal variations in O 2 at the surface are not expected to propagate to the upper atmosphere. Recent measurements made through disk‐averaged observations of Mars with the Herschel Space Observatory's HIFI instrument retrieved a value of 1400 ± 120 ppmv, though they caution the reader the value may not be vertically uniform (Hartogh, Jarchow, et al., 2010). This measurement was taken during MY 30, L S 77°. SAM measurements from surface from similar times of year are slightly higher but with overlapping uncertainty (Table S1).

2 in Gale Crater do not show the annual stability or seasonal patterns that would be predicted based on the known sources and sinks in the atmosphere. As mentioned in section 2 should show the same seasonal patterns and annual repeatability as Ar. Given the known chemical cycles, the formation of atmospheric O 2 is controlled primarily by photochemistry of H 2 O and CO 2 (e.g., Atreya & Gu, 1994 (R1) (R2) (R3) (R4) 2 ). The abundance is then controlled by the balance between formation and loss through photolysis and formation of H 2 O 2 , HO 2 , and NO 2 (Krasnopolsky, 1993 The SAM measurements of Oin Gale Crater do not show the annual stability or seasonal patterns that would be predicted based on the known sources and sinks in the atmosphere. As mentioned in section 3.2.2 , based on known sources and sinks, Oshould show the same seasonal patterns and annual repeatability as Ar. Given the known chemical cycles, the formation of atmospheric Ois controlled primarily by photochemistry of HO and CO(e.g., Atreya & Gu,):andwhere M is the background gas (CO). The abundance is then controlled by the balance between formation and loss through photolysis and formation of H, HO, and NO(Krasnopolsky,).

To quantify the enrichment of the observed O 2 compared to what would be predicted, we account for the changes in total number density, which are mainly due to pressure changes (Figure 7a) caused by the global‐scale CO 2 condensation‐sublimation cycle, as well as the expected changes in the mixing ratios of all noncondensable gases, which are caused by the interaction of the global circulation with that condensation‐sublimation cycle and are readily visible to us via Ar results. After accounting for both of these, it can be estimated that approximately 1014 O 2 molecules cm‐3 must be added to the atmosphere sampled by Curiosity in order to explain the observed 1,700 to 2,200 ppmv increase in O 2 between L S 60°and 140° in MY 33. In other words, the number density of O 2 molecules at L S 140° in MY 33 is ~1014 molecules cm‐3 larger than it would have been had the ratio of O 2 to Ar remained constant as expected. For reference, given the 1,700 ppm starting value at L S 60°, the O 2 mixing ratio would have been ~1830 ppm at L S 140°, had the O 2 to Ar ratio remained constant over that time span.

Using the O 2 /40Ar ratio as an indicator for O 2 variability outside of the known and observed seasonal dynamics, a simple linear model can be used to fit a single slope to all the data (Figure 10). This model highlights a consistent seasonal increase in O 2 /40Ar ratio of 0.014/100° L S during the L S 0–150° period, with an interannual variation in the mean O 2 /40Ar ratio in this period. For L S >150°, O 2 /40Ar seems to be more or less constant, with no significant interannual variation. The low values at L S >310° in MY 31 suggest a possibility of additional interannual variation late in the Mars year, and potentially the onset of the increases is observed through the spring of the following years, but future observations would be needed to confirm this possibility.

Within the uncertainties caused by limited sampling and measurement error, this magnitude appears typical of the unexpected seasonal increase, and so going forward, we will adopt ~400 ppm and ~014 molecules cm‐3 as the amount that needs to be resolved. Assuming that the unexpected O 2 is uniformly mixed in the lower atmosphere, as seems likely for perturbations of this timescale given current assumptions about the eddy diffusion coefficient, the 1014 molecules cm‐3 become 1020 molecules cm‐2 in the atmospheric column (see, e.g., Krasnopolsky, 2010 who adopts 107 cm2 s‐1 for the eddy diffusion coefficient, which gives a ~2‐day characteristic timescale for the bottom scale height of the atmosphere).

Given photochemical schemes above, this 400 ppm of extra O 2 would require a corresponding destruction of CO 2 and H 2 O molecules in approximately 170 sol. Considering H 2 O alone, ~800 ppm of H 2 O would need to be destroyed, which is more than five times larger than the maximum abundance of H 2 O measured in and around Gale Crater by REMS (Martínez et al., 2016) and ChemCam (Fig. 111 in McConnochie et al., 2018). Thus, it appears unlikely that the needed O 2 could be produced from the available atmospheric water for any plausible H 2 O photolysis or dissociation mechanism. Furthermore, H 2 O abundance shows an increase during this time period and no strong correlation with O 2 (Figure S6). Estimates for the production of O from CO 2 , using CO 2 photolysis rates for the lower atmosphere of Mars (Table 1 in Wong et al., 2003), indicate that this process is much too slow to generate the observed rise over the short (~1/2 yr) time period. For completeness, note that photolysis or other dissociation of CO is negligible, and in any case a seasonal removal of ~800 ppm of CO is clearly ruled out by observations (e.g., Smith et al., 2009).

The primary destruction pathways for O 2 are through direct photolysis in the upper atmosphere and reaction with photolysis products of H 2 O, HO 2 , and CO 2 deeper (Atreya et al., 2006; Lefèvre & Krasnopolsky, 2017; Wong et al., 2003). Even factoring in the effects of dust devils and large dust storms (though no major dust events occurred during the measurement period) (Atreya et al., 2006), the lifetime of O 2 against photochemical destruction in the Mars atmosphere is expected to be at least 10 years, possibly longer (Krasnopolsky, 2010; Lefèvre & Krasnopolsky, 2017). Again using the O 2 /40Ar ratio as an indicator for O 2 variability (Figure 10), the SAM measurements in MY 33 show a relative decrease of 23% in a period of 39 days (38 sol). (This is a ~500 ppm absolute O 2 decrease from what would be expected from a constant O 2 /40Ar ratio given the starting O2 abundance at L S 140°.) There is a similar decrease observed from fall to winter of MY 31, although the measurement frequency is such that the period of change appears longer (–20% in 201 sol). In particular, the rapid drop in MY 33 corresponds to a lifetime of 150 days, several orders of magnitude less than the photochemical equilibrium lifetime of ≥10 years. The drop in MY 31 corresponds to a lifetime of ~1000 days, which is still relatively short.

The lack of a known atmospheric source or sink that could explain the apparent behavior of the O 2 in Gale Crater suggests the possibility of a temporary surface reservoir. Previously, Herschel WIFI observations found that the O 2 vertical profile above the surface is not constant with altitude (Hartogh, Jarchow, et al., 2010) and preliminary analysis of the data shows that a surface flux of O 2 may be required to explain the observations (Paul Hartogh, personal communication, 2016). A surface sink has been previously invoked to balance the current redox budget (Zahnle et al., 2008), and the surface is known to harbor a variety of oxidant species (e.g., Lasne et al., 2016 review). In fact, the Viking Gas Exchange experiments found that a significant quantity of O 2 was released whenever soil samples were humidified (Klein, 1978; Oyama & Berdahl, 1977) although these experiments were all done at ~10 °C rather than Mars ambient temperatures.

Deposition of oxygen could occur in the form of more reactive oxidized species, superoxides, hydrogen peroxide (H 2 O 2 ), ozone (O 3 ), or perchlorates, all of which are assumed to have a higher surface reaction probability (γ) with surface materials than molecular oxygen. It is possible to conceive of an oxygen cycle with the appropriate seasonal and interannual variability if oxygen were effectively converted to these species, deposited into the regolith, and then rereleased due to thermal, chemical, or radiation perturbations.

Perchlorates, found to be prevalent in the surface materials in Gale Crater at 0.03–1 wt% level (Sutter et al., 2017), are very stable. They have also been detected at 0.4–0.6 wt% level in the polar landing site of the Phoenix Lander (Hecht et al., 2009). To put perchlorates in context, a 1‐cm depth of soil containing 1% by weight of calcium perchlorate has slightly more than enough oxygen to contribute the apparent ~1020 molecules cm‐2 of unexpected column O 2 variation, and O 2 has been shown to be a high yield product of radiolysis of surface perchlorate salts (Quinn et al., 2013). However these results point to a long‐term accumulation of O 2 in the Martian soil; the production rate of O 2 from perchlorate radiolysis is insufficient to produce the observed recurring unexplained signal. More specifically, based on the Quinn et al. (2013) experimental radiation dose and yield, and on their estimated Martian dose rates and estimated 2‐m cosmic ray penetration depth, it would take on order of 1 Myr to accumulate enough trapped O 2 for one season worth of 1020 molecules cm‐2 of column O 2 variation. Similarly, the proposed “superoxide” O 2 ‐ ions, suggested to explain the results of the Viking soil reactivity experiments, could form from ultraviolet radiation on surface minerals and lead to the observed release of O 2 with humidification (Yen et al., 2000), but the reported rate of superoxide generation is too small to be consistent with the inferred column O 2 signal yet again by a factor of ~106.

The Viking Gas Exchange experiments released up to 770 nanomoles of O 2 from a 1 cm3 sample over a period of less than 11 days upon “humidification” at ~10° C (Klein, 1978; Oyama & Berdahl, 1977). If the same abundance of rapidly releasable O 2 was present across 2 m of depth (i.e., 200 cm3), this would yield the 1020 molecules cm‐2 that would explain the atmospheric measurements. This indicates that sufficient rapidly releasable O 2 is present in the Martian soil, although it is not clear that such a rapid release of O 2 could have occurred at Mars ambient temperatures. More importantly, this serves to illustrate that explaining a one‐time release of O 2 is not the main problem. The primary difficulty is that the slow rates of accumulation in the processes considered so far cannot explain the seasonal recurrence of excess O 2 .

Hydrogen peroxide is worth considering as a solution to this problem, because it is less stable than perchlorates, and is therefore more likely to provide a rapidly exchangeable reservoir of O 2 . H 2 O 2 has been detected in the atmosphere (Clancy et al., 2004; Encrenaz et al., 2004) and exhibits seasonal and interannual variability (Encrenaz et al., 2019), and it has been suggested that diffusion of atmospheric H 2 O 2 into the regolith and/or mineral/water interactions could supply H 2 O 2 in the near subsurface (Bullock et al., 1994; Lasne et al., 2016).

Current knowledge of H 2 O 2 physics and chemistry in Martian soil is very limited, but based on a coupled soil‐atmosphere model by Bullock et al. (1994), it appears that the magnitude and thermal sensitivity of H 2 O 2 soil adsorption is potentially close to the right order of magnitude to supply the unexpected 1020 molecules cm‐2 of O 2 . However, this conclusion only follows from assuming a 107‐year chemical lifetime for adsorbed H 2 O 2 , which is the longest that Bullock et al. (1994) considered plausible, and it depends on their adopted absorption isotherm, which was based on the Fanale and Cannon (1971) empirically derived expression for H 2 O since no data for H 2 O 2 were available. Furthermore, once the depth penetration of the annual temperature wave (e.g., Grott et al., 2007) is considered, the amount of H 2 O 2 potentially cycled in and out of the soil is estimated at an order of magnitude less than needed here, at 1019 molecules cm‐2. Further, the timescale for diffusion from meter‐scale depths may be far too long. Finally, note that although seasonal trends are in fact observed for atmospheric H 2 O 2 , this mechanism would require a rapid conversion to O 2 immediately at the surface as the observed amount of H 2 O 2 is only on the order of ppb (Encrenaz et al., 2015).

Another potential mechanism involving H 2 O 2 was proposed by Quinn and Zent (1999). They showed that H 2 O 2 ‐TiO 2 complexes rapidly released O 2 upon humidification at warm (for Mars) temperatures of 10 °C; a more extensive study of regolith analogs and environmental conditions for this type of H 2 O 2 ➔ O 2 conversion in the regolith is needed.