Effect of geologically-constrained environmental parameters on the

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Effect of geologically-constrained environmental parameters on the atmosphere and biosphere of early Earth Stefanie Gebauer, John Lee Grenfell, Ralph Lehmann, and Heike Rauer ACS Earth Space Chem., Just Accepted Manuscript • DOI: 10.1021/ acsearthspacechem.8b00088 • Publication Date (Web): 21 Sep 2018 Downloaded from http://pubs.acs.org on September 25, 2018

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ACS Earth and Space Chemistry

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Effect of geologically-constrained environmental parameters on the atmosphere and biosphere of early Earth

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S. Gebauer1,? , J. L. Grenfell1 , R. Lehmann2 , and H. Rauer1,3

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(1) Institut für Planetenforschung (PF), Abteilung Extrasolare Planeten und Atmosphären (EPA), Deutsches Zentrum für Luft- und Raumfahrt (DLR), Rutherfordstr. 2, 12489 Berlin, Germany (2) Alfred-Wegener Institut, Helmholtz-Zentrum für Polar- und Meeresforschung, Telegrafenberg A 45, 14473 Potsdam, Germany (3) Zentrum für Astronomie und Astrophysik (ZAA), Technische Universität Berlin (TUB), Hardenbergstr. 36, 10623 Berlin, Germany

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Running title: Geological constraints affecting early Earth

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ACS Earth and Space Chemistry 1 2 3 4 5 6 7 8 9 10 11 12 13 14 15 16 17 18 19 20 21 22 23 24 25 26 27 28 29 30 31 32 33 34 35 36 37 38 39 40 41 42 43 44 45 46 47 48 49 50 51 52 53 54 55 56 57 58 59 60

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Corresponding author: Dr. Stefanie Gebauer Institut für Planetenforschung (PF) Abteilung Extrasolare Planeten und Atmosphären (EPA) Deutsches Zentrum für Luft- und Raumfahrt (DLR) Rutherfordstr. 2 12489 Berlin Germany phone: +49-(0)30-67055-454 e-mail: [email protected]


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Abstract

The Great Oxidation Event (GOE) about 2.3 Gigayears ago denotes the first major rise of atmospheric molecular oxygen (O2 ) in Earth’s history. As a consequence the planet experienced the emergence of widespread habitability and complex life. Recently there has been a revolution in improved methods for constraining geological data for atmospheric pressure, composition and ocean temperature of the early Earth. We investigate the effect of this revised data upon processes which drove the GOE. Results suggest that increasing Archean carbon dioxide (CO2 ) produces increased O2 with height due to enhanced CO2 photolysis. This is counterbalanced by stronger O2 destruction as a result of enhanced carbon monoxide (CO) (from increased CO2 ) and nitrogen oxides (NOx ) (from decreased hydroxyl via increased CO). Pre-GOE atmospheres with low O2 yet high CO2 could counteract O2 accumulation. For low surface pressures of 0.5 bar, O2 decreases between 0.5 to 0.005 bar. This arose mainly from O2 destruction via hydrogen oxides from enhanced water from higher temperatures for p < 0.01 bar and weaker O2 production via less ultraviolet radiation that initiates ozone production via CO2 photolysis. Shortly before the GOE, ∼ 20% lower Net Primary Productivity (NPP) can maintain comparable O2 as for a 1 bar atmosphere and, hence, the accumulation of O2 produced by a photosynthetic biosphere is supported. We identify new O2 production and destruction pathways with NOx containing species for Archean Earth for high CO2 atmospheres and low/high surface pressure. On increasing ocean temperatures, NPP is reduced due to lower O2 solubility before the GOE. This facilitates atmospheric O2 accumulation. Key words: early Earth – atmosphere – biogeochemistry – proxy data – Oxygen – Carbon dioxide



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Introduction

The Great Oxidation Event (GOE) which likely began about 2.4 Gigayears (Gyrs) ago marked a spectacular rise of molecular oxygen (O2 ) in the atmosphere - from less than 10−6 of the Present Atmospheric Level (PAL) 1 up to at least 0.01 PAL 2–4 in the Paleoproterozoic. This event enabled wide-ranging habitability due to the formation of a significant ozone (O3 ) layer which protects life on the surface from Ultraviolet (UV) radiation and also allowed complex life to take hold of the planet since this requires an oxygen-rich atmosphere for obtaining energy via respiration. Oxygenation led to a lethal impact upon methanogens and there followed a series of ice age events 5–7 marking the Neoproterozoic illustrates the extremity of this event. Interplays within the atmosphere-biology-geology system are therefore both complex and far-reaching. Many questions still remain open regarding the underlying mechanism which drove the GOE, many questions still remain open. Several theories have been proposed e.g. due to a stronger overall O2 source associated with an increase in the burial rate of organic carbon 8–10 which was possibly related to the growth of continental shelves or via enhanced burial efficiency due to strong hydrothermal activity 11 . However, this latter point was contested 12 . An alternative explanation for the GOE is due to a weaker O2 sink possibly associated with a decline in the reducing nature of volcanic gases via a change in the oxidation state of the mantle 13 . Some authors 14 have however disputed a step-wise reduction in mantle oxidation based on an analysis of redox-sensitive elements (e.g. chromium) in basalt rock. The GOE was also proposed to be related to the long-term escape of hydrogen 3 . Applying proxy data to constrain the timing and extent of the GOE is a rapidly expanding field. Sulphur-Mass Independent Fractionation (MIF-S) data was suggested 15 to indicate strong surface photolysis of sulfur dioxide (SO2 ) which could provide indirect evidence of strong (near) surface UV due to low O3 (and, hence, low O2 ) in the atmosphere prior to ∼ 2.4 Gyrs ago. The precise mechanism linking processes in the atmosphere with the origin of MIF-S in the rock record is however still unclear 16 . There is also evidence of the presence of locally enhanced amounts of O2 even prior to the GOE 17 from low MIF-S, redox-sensitive trace element studies 18–20 and isotope data analysis 21 . This could be evidence for a so-called ”yo-yo” behavior (as opposed to a linear rise) in O2 during the GOE due to the proposed oxygen bistability 22 although this point is contested 23 . Interpreting MIF-S could also have repercussions for carbon dioxide (CO2 ), methane (CH4 ), molecular hydrogen (H2 ) or/and aerosols in the atmosphere since these species could have impacted surface UV or/and sediment formation rates 24 . Molybdenum isotope data 25 from sedimentary rocks support the theory that


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the surface ocean experienced oxygenation during the GOE yet the deep ocean remained mostly reducing up until the Second Oxidation Event (SOE) which started about 600 million years ago 4,26 . Banded Iron Formations (BIFs) indicate a fluctuation in the oxidation state of the atmosphere-ocean system. The co-existence of magnetite (an oxidized mineral) with siderate (a reduced mineral) in BIFs depended on the abundance of atmospheric H2 27 . Since a minimum amount of H2 is required for methanogenic respiration, they therefore derived a minimum partial pressure of atmospheric CO2 of 3 PAL although their conclusions are contested 28 . Other studies 29–32 which considered paleosols have indicated larger amounts of CO2 up to several hundreds of PAL. There are still open questions regarding our knowledge of atmospheric processes affecting habitability on the early Earth. The Faint Young Sun Problem (FYSP) refers to the contradiction of liquid water likely existing on Early Earth’s surface as evidenced by e.g. detrital zircon data 33,34 despite the fainter early Sun 35 . In order to address the FYSP, numerous model studies suggested greenhouse gases such as CO2 36–38 or CH4 39–41 . Amounts required however, are in conflict with geochemical constraints derived from paleosols (for references see above) or are problematic because they would promote cooling instead of warming due to the generation of hazes 40,42 . However, it was recently shown 43 that organic hazes are self-limiting due to their self-shielding properties, preventing extreme cooling of a planet. Other greenhouse gases such as ethane 40 , nitrous oxide 44 and carbonyl sulfide 45 have also been proposed. Furthermore, the greenhouse effect could have been amplified by pressure broadening of absorption lines due to enhanced pressure by an increase in molecular nitrogen (N2 ) 46 . In summary, uncertainties in the radiative transfer functions as well as the lack of spatially-resolved and fully coupled climate models for the early Earth including the full range of feedbacks in the Earth system, makes a final assessment of greenhouse-gas warming in the early atmosphere challenging. Other climatic factors such as changes in cloud cover and surface albedo could at least have contributed to a warming of the Archean Earth. Numerous 3D model studies investigating the FYSP have appeared. These have suggested a continuously warm equatorial zone 47,48 consistent with the sediment data. 3D cloud feedbacks could also have played an important role 49 . The enhanced greenhouse warming due to H2 and N2 could also address the FYSP 50 Some recent studies 51 investigated possible glaciation events potentially associated with the GOE. For a comprehensive review see 52 and references therein. Understanding the GOE is clearly linked with understanding the FYSP, since life (and, hence, ultimately the rise in O2 ) requires habitable conditions. Implications for the surface pressure of the early Earth have been derived from fossilized raindrops 53 . They suggested rather modest surface pressures, psurf , i.e.



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below 2 bar (likely 1.1 bar), at ∼2.7 Gyrs, although the raindrop barometry method is contested 54 . Also, N2 amounts between 0.5 to 1.1 bar during the mid-Archean based on a nitrogen isotope analysis of hydrothermal quartz have been proposed 55 . In a more recent study 56 they performed an analysis of the size distribution of gas bubbles in basaltic lava flows that solidified at sea level ∼2.7 Gyr in the Pilbara Craton, Australia. Their study suggested an even lower surface pressure of psurf = 0.23 bar and an uppermost value of 0.5 bar. Since geochemical data point towards temperatures above the freezing point of water, the question arises as to how warm the Archean climate actually was. Several studies point towards either a hot (∼70◦ C) 57–60 or a more temperate ocean 61,62 3.5 Gyrs ago where the former is highly debatable 17,63–65 . A recent study 66 reviews different interpretations of oxygen and silicon isotope variation and applies a model of formation based on ancient and modern chert studies yielding intermediate temperatures between (37 - 52)◦ C for the Precambrian ocean. A recent study 67 also provides additional arguments for a temperate Archean climate (0 - 50) ◦ C derived by applying a geological carbon cycle model with ocean chemistry.

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Contribution of this Work

The evolution of early Earth analog planets around the Sun and the implications for their atmospheres and biospheres was addressed 68 (hereafter paper I) by applying a Coupled Atmosphere Biogeochemical (CAB) model with its detailed oxygen cycle components (e.g. burial, weathering, photochemistry) from the atmosphere, biosphere and geosphere. In this work we apply the CAB model (described in detail in paper I) to early Earth scenarios based on recent advances in geological data for surface pressure, atmospheric greenhouse gases and ocean temperature. This model is able to derive the corresponding net primary productivity of a photosynthetic biosphere which is needed to maintain a specified O2 concentration at the surface. Hence we are able to address the impact of changing surface pressure, atmospheric greenhouse gases and ocean temperature upon the O2 producing biosphere. In order to simulate these new scenarios we have introduced major updates into the climate and photochemistry modules of the CAB model. With the updated model we are able to investigate the interplay of key processes which drove the GOE in the atmosphere of early Earth. To understand the atmospheric chemical responses of O2 in detail for the different Archean scenarios considered we apply the Pathways Analysis Program (PAP) 69 . The paper is structured as follows: Section 3 and Section 4 describe briefly the models used and their updates. Section 5 is the CAB model validation for modern Earth and comparison with the previous model version followed by Section


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6 which gives a brief overview of the simulated early Earth scenarios. Sections 7, 8 and 9 present the results for varying the CO2 content, surface pressure and ocean temperature, respectively. Section 10 shows the main conclusions.

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Coupled Atmosphere Biogeochemical (CAB) Model

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The climate and photochemistry modules of the CAB model were updated in order to address the interactions between atmospheric, geological, chemical and biotic systems. A more detailed description of the CAB model is provided in paper I.

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Climate module updates

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In order to simulate the full range of atmospheres (pCO2 , surface pressure etc.) implied by the latest geological dataset as described above, the existing LongWave (LW) radiation scheme of the CAB model (namely the RRTM (Rapid Radiative Transfer Module) 70 scheme) was replaced by the more flexible MRAC-complete (Modified RRTM for Application in CO2 -dominated Atmospheres) LW scheme 71 and based on the original MRAC scheme 38,72 . The MRAC-complete scheme is originally based on the radiation scheme RRTM. It considers the absorption by radiative gases CO2 , H2 O, CH4 and O3 as well as the self- and foreign-continua for CO2 and H2 O. It can however be applied over a much wider range of surface pressure (up to psurf = 1.5 bar), H2 O (10−9 − 1), CO2 (10−6 − 1), CH4 (10−8 − 10−2 ) and O3 (10−8 − 10−2 ) volume mixing ratios as well as temperatures (up to T = 700 K) compared with the existing RRTM scheme. The MRAC scheme operates from 1 to 500 microns split into 25 bands and uses the correlated-k method (see e.g. von Paris et al., 2008). The pressure and temperature boundaries are limits of the LW radiation module and arose when calculating the correlated-k coefficients. A more flexible model version with a wider pressure and temperature range is under development. Further updates include heat capacity parameterizations for H2 O, CH4 , H2 , He and CO and Rayleigh scattering parameterizations for these 5 species in addition to the existing parameterizations for O2 , N2 and CO2 in the CAB model. For a more comprehensive description we refer to 71 . Furthermore, in MRAC-complete the pressure grid is now adjusted according to the changing mass of the atmosphere due to the evaporation of both H2 O and CO2 (not only H2 O as before).

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3.2 Photochemistry updates

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Photolysis in the middle to upper atmosphere is a major atmospheric sink for O2 and is important for understanding its atmospheric budget. It is a key process forming



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atomic oxygen (O) which is required for forming the O3 layer and, hence, protecting the surface from harmful UV radiation. In the course of this work the existing parameterization for the O2 Schumann-Runge (S-R) bands (between 175 and 193 nm) were updated. Without this update the O2 photolysis rate can be underestimated by two orders of magnitude 73 . We have implemented the effective O2 column density dependent O2 cross sections 74 for each respective S-R band of the photochemistry module. In order to validate the updated CAB model to reproduce modern Earth we changed the zenith angle from 45◦ to 53◦ since with 45◦ the O3 column amount was overestimated due to the new increased O2 photolysis rate in the S-R wavelength range (see also discussion on appropriate mean zenith angle for the application in 1-dimensional photochemical models 75 ). The new zenith angle reproduces the mean O3 column amount for modern Earth 76 . Furthermore, Rayleigh scattering parameterizations for H2 O, CH4 , H2 , He and CO were implemented in order to be consistent with the climate module.

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Pathways Analysis Program

We apply the Pathway Analysis Program (PAP) 69 to automatically identify chemical pathways in any arbitrary chemical reaction network and to quantify their efficiencies by assigning pathway rates. PAP has been applied in various studies 68,77–85 . All dominant pathways which produce, destroy or recycle a chemical species of interest are found by PAP. For a more detailed description of PAP see paper I and 69 . Since the number of chemical pathways increases rapidly with the size of the chemical reaction network, it is necessary to identify and discard unimportant pathways, i.e. chemical pathways with a relatively low throughput flux, in an early stage of the analysis. Therefore a threshold rate fmin is set by the user. All pathways with a rate smaller than this threshold rate are deleted. If fmin is chosen to be too large, potentially important pathways will be deleted and the total production or consumption of the species of interest can not be completely explained by the pathways determined using the algorithm. On the other hand, if fmin is chosen to be too small, the required computational time and memory consumption increase rapidly. In the present study, fmin = 10−13 parts per billion by volume per second (ppbv/s) was chosen to be sufficient for finding the dominant O2 production and destruction pathways in our simulated early Earth atmospheres. This value reflects the results of previous studies such as Stock et al. (2012b). A PAP analysis was performed for each of the 64 vertical layers in the photochemical module. The resulting production and destruction rates of O2 from each individual pathway are integrated over the vertical grid of the atmospheric module and are expressed as a percentage of the total column-integrated production and destruction rates of all pathways found by


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CAB Model Validation and Comparison with Previous Model Version

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The updated CAB model i.e. with the module ”MRAC-complete” replacing ”RRTM” and with the updated S-R O2 scheme is hereafter referred to as CABMRACO2. It was applied to validate the model results for modern and Archean Earth and compared with the results of paper I (denoted as CABRRTM) and a CAB model version with MRAC-complete but without the updated O2 photolysis scheme (denoted as CABMRAC). CABRRTM was validated against modern Earth and reproduced modern Earth surface temperature of 288.15 K as well as observed surface mixing ratios of the chemical species H2 = 5.5 · 10−7 , CH4 = 1.6 · 10−6 , N2 O= 3 · 10−7 , CO= 9 · 10−8 , and CH3 Cl= 5 · 10−10 . As in 76 we are able to reproduce the main features of the 1976 U.S. Standard Atmosphere temperature profile, O3 , and water profiles. In order to achieve a surface temperature of 288.15 K the surface albedo was adjusted (from 0.2067 in paper I to 0.212 in the present work) as was done for previous model studies 76,86 . The climate and photochemistry analyses are shown in detail in Appendix A.

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6 Scenarios for Early Earth Analog Atmospheres

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Tab. 1: Early Earth scenarios assumed in this work. The values of the geological time, tgeo , in column 2 are determined such that the corresponding surface O2 volume mixing ratio (vmr) in column 3 coincides with the possible evolutionary path 3 of O2 . eon Proterozoic

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In order to analyze the impact of a variation in geologically constrained parameters such as CO2 content, surface pressure and ocean temperature upon the atmosphere



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and biosphere of early Earth, the CABMRACO2 model is applied to early Earth scenarios, namely before, during and after the GOE. These are described in Tab. 1. For each scenario we moderately vary either the CO2 content (1 PAL = 355 ppm, 10, 100 PAL) or the surface pressure (psurf = p0 +pH2 O +pCO2 with the background pressure p0 = 0.5, 1, 1.5 bar) to cover the range of proposed values in the literature presented in Section 1. pH2 O and pCO2 denote the partial pressure of H2 O and CO2 , respectively. The O2 volume mixing ratio of the atmosphere at the surface is taken from the evolutionary path 3 of O2 . Note, that this pathway was chosen only as an illustrative model and that there is great uncertainty in Proterozoic O2 values. Our set of scenarios neglects a possible O2 overshoot 87 . All early Earth scenarios use the stellar spectrum of the Sun with the total incoming (top-of-atmosphere) flux scaled with age 35 . Furthermore, for all scenarios we assume: • gravitional acceleration, g = 981 cm s−2 , • biogenic surface fluxes of CH4 (360 Tg/yr), CO (1476 Tg/yr), N2 O (14.8 Tg/yr), CH3 Cl (2.7 Tg/yr) (these values lead to modern Earth abundances for the given species), H2 deposition velocity (vdep = 2.54 · 10−3 cm s−1 ), deposition velocities of all other chemical compounds as in paper I, • crustal mineral redox buffer is set to Quartz-Fayalite-Magnetite (QFM) (∆fO2 = 0). Volcanic emissions of H2 , H2 S, SO2 , CO, CH4 , and CO2 and metamorphic emissions of H2 and CH4 depend on the heat flow parameter Q, which diminishes on going forward in geological time tgeo , and are varied as in paper I. The metamorphic outgassing depends on the assumed crustal mineral redox buffer. In order to investigate the effect of varying ocean temperature, Tocean , upon the NPP from oxygenic photosynthesis, we vary the ocean temperature, Tocean , from 25 ◦ C (default value as in paper I) to 70 ◦ C as proposed in the literature for a hot ocean (see Section 1). The piston velocity (gas exchange rate across the oceanatmosphere boundary) changes 88 from vp (O2 ) ≈ 5.7 · 10−3 cm s−1 at 25 ◦ C to vp (O2 ) ≈ 1.4 · 10−2 cm s−1 at 70 ◦ C. The solubility constant reduces from HO2 = 0.0013 M/atm 89 at 25 ◦ C to HO2 = 0.00068 M/atm 90 at 70 ◦ C.


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Fig. 1: Selected p−T profiles of early Earth analog atmosphere scenarios calculated with the fully updated CAB model (CABMRACO2). 288

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Sensitivity with Respect to Atmospheric CO2 Concentration Atmospheric Modeling

Fig. 1 presents p-T profiles for selected early Earth atmosphere scenarios before, during and after the GOE for varying atmospheric CO2 content and p0 = 1 bar. A temperature inversion becomes visible with increasing O2 content for 1 PAL CO2 (as was also the case in paper I) since increased O2 favors the formation of O3 which heats the middle atmosphere. On increasing the CO2 content to 100 PAL the temperature inversion diminishes due to stratospheric cooling by CO2 whereas surface temperatures increase as expected due to greenhouse warming by CO2 . In the case of the Archean scenarios with 10−6 PAL O2 and 10 and 100 PAL CO2 stratospheric temperatures increase towards lower pressures because of heating by enhanced O3 formed from O2 due to increased CO2 photolysis. On going forwards in time, i.e. increasing O2 surface vmr, surface temperatures (see Fig. 2) increase from 265, 272 and 282 K before the GOE at 10−6 PAL O2 to 288, 296 and 307 K at 1 PAL O2 for 1, 10 and 100 PAL CO2 respectively. Hence, to satisfy geochemical constraints from zircon data 33,34 which indicate the presence of



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Fig. 2: Surface temperature evolution of early Earth analog atmosphere scenarios calculated with the fully updated CAB model (CABMRACO2). Additionally indicated as a horizontal line is the freezing temperature of H2 O of T = 273.15 K.

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liquid water on the surface of early Earth 4.4 Gyrs ago, our results suggest that the CO2 vmr must have been higher than 10 PAL CO2 . A previous study 38 calculates a minimum surface CO2 partial pressure of about 10 mbar (see their Fig. 11) for a CO2 -H2 O atmosphere at a solar constant of S = 0.81 (equivalent to our scenarios with 10−6 PAL O2 ). Compared to that work, we additionally consider the radiative properties of O3 and CH4 (MRAC-complete) as well as coupled photochemistry. We require less CO2 i.e. about 6 mbar to counteract the effect of the faint young Sun and maintain habitable surface conditions. Note that in the case of our scenarios prior to 1.9 and 2.5 Gyrs ago 1 and 10 PAL CO2 , respectively, global surface temperatures are below 0◦ C implying no habitable surface conditions. However, it was shown by 3D general circulation models investigating the early Earth 47,48 that even with global mean surface temperatures below the freezing point of water, low latitude regions with liquid surface ocean water could still persist implying habitable conditions.

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7.1.2 Photochemical Responses

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UV Radiation environment Fig. 3 shows the UVA, UVB and UVC radiation profiles for the same scenarios as in Fig. 1. Note that the x-axis differs for each panel. UVA radiation is absorbed efficiently such that fluxes are lowered by about 40% compared to values at the Top-Of-Atmosphere (TOA) in all calculated scenarios and no major differences are evident on varying the atmospheric CO2 content. UVB, however, is absorbed in the middle atmosphere both during and after the GOE whereas before the GOE UVB radiation is absorbed only in the lower atmosphere. On increasing CO2 content, the absorption of UVB radiation is enhanced due to more O3 formed from pathways initiated by CO2 photolysis (see Fig. 5). This


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Fig. 3: UVA (top), UVB (middle) and UVC (bottom) radiation profiles in W/m2 of early Earth analog atmospheric scenarios calculated with the fully updated CAB model (CABMRACO2).



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Fig. 4: Ratio (surface/TOA) radiation fluxes, RUV , for UVA (black), UVB (red) and UVC (green) of early Earth analog atmosphere scenarios calculated with the fully updated CAB model (CABMRACO2) as a function of geological time tgeo . 329 330 331 332 333 334 335 336 337 338 339 340 341

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effect is even stronger for UVC radiation before the GOE where it is efficiently absorbed (also due to enhanced H2 O in these atmospheres) although a few W per m2 still reach the surface. During and after the GOE, surface UVC is virtually zero. The radiation shielding factor, RU V , (i.e. the ratio (surface/TOA) radiation) is shown in Fig. 4 of such atmospheres. Low RU V values denote high shielding efficiency. All values are reduced in comparison to paper I and, hence, UV radiation is blocked more efficiently due to enhanced O3 and H2 O. On increasing CO2 , atmospheres become more wet due to stronger evaporation for the higher surface temperatures. Moreover, there is enhanced O3 production initiated by CO2 photolysis (see analysis below). Therefore, the effect of increasing CO2 is strongest in the case of UVC before the GOE. In the context of early Earth we will now focus on the photochemical analysis of O2 and O3 . Ozone - O3 Fig. 5 presents selected O3 profiles of early Earth atmospheric scenarios before (blue lines), during (red lines) and after (black lines) the GOE. The impact of enhancing CO2 is strongest in the high UV environment before the GOE due to pathways initiated by CO2 photolysis producing O2 and subsequently forming O3 . In the case of the Archean scenario with 10−6 PAL O2 and 100 PAL CO2 the O3 concentrations exceed that of the atmospheric scenarios with 10−3 PAL O2 and 10 PAL CO2 in the upper atmosphere. In the lower atmosphere before the GOE O3 concentrations increase with increasing CO2 whereas for the scenarios during and after the GOE the reverse is the case due to decreased O and O(1 D) concentrations in the lower atmosphere during and after the GOE. This occurs due to a weakening in the UV environment in the lower atmosphere associated with increased UV shielding by the now-established O3 layer during and after the GOE.


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Fig. 5: Selected O3 profiles of early Earth analog atmosphere scenarios calculated with the fully updated CAB model (CABMRACO2). 354 355 356 357 358 359 360 361 362

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The evolution of the O3 column amount for different atmospheric CO2 vmrs is given in Fig. 6. In order to visualize the impact of increasing CO2 amount before and after the GOE we show here two panels with different y-axes, i.e. linear (upper) and logarithmic (lower). Before the GOE the O3 column amount increases on increasing the CO2 content contributing to the formation of O3 which can help to shield the surface from UV radiation. However, after the GOE the largest O3 column amount arises for an atmosphere with 10 PAL CO2 . This is related to enhanced O3 concentrations in the upper atmosphere with increasing CO2 but a stronger decrease of O3 in the lower atmosphere for the 100 PAL CO2 scenario. Molecular oxygen - O2 Fig. 7 shows selected O2 profiles for the calculated early Earth scenarios. Before the GOE, O2 is strongly increased on increasing CO2 whereas this is not the case for the scenarios during and after the GOE because the number of produced O2 molecules is small relative to the total available O2 . Also weak UV shielding before the GOE implies that the role of CO2 is more important. In the case of 100 PAL CO2 one obtains O2 concentrations (below 10−3 bar) which are close in value to the scenarios during the GOE. For a detailed chemical pathway analysis of O2 sources and sinks see below. Fig. 8 shows the evolution of the atmospheric O2 surface flux which is needed to



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Fig. 6: Evolution of O3 column [DU] amount of early Earth analog atmosphere scenarios calculated with the fully updated CAB model (CABMRACO2). Top: linear y-axis. Bottom: logarithmic y-axis.


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Fig. 7: Selected O2 profiles of early Earth analog atmosphere scenarios calculated with the fully updated CAB model (CABMRACO2).

Fig. 8: Evolution of O2 surface flux [Tg/yr] of early Earth analog atmosphere scenarios calculated with the fully updated CAB model (CABMRACO2).



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Fig. 9: O2 net chemical change (P −L) for early Earth analog atmosphere scenarios with a surface O2 vmr of 10−6 PAL O2 calculated with the fully updated CAB model (CABMRACO2) for 1 PAL CO2 (solid) and 100 PAL CO2 (dotted). The vertical dot-dashed line indicates the zero line. 372 373 374 375 376 377 378 379 380 381 382 383 384 385 386 387 388 389 390 391 392 393

maintain a given O2 surface mixing ratio (taken from proxy data 3 ) for 1, 10 and 100 PAL CO2 . The O2 surface flux for the 100 PAL CO2 atmosphere is highest before the GOE. This indicates that the net (P − L) chemical change is more negative (cf. Fig. 9) and, hence, these atmospheres are chemically more destructive towards O2 (although O2 production from CO2 photolysis is higher). After the GOE, higher amounts of CO2 in the atmosphere result in lower net chemical O2 destruction. In the case of the scenario with 100 PAL CO2 , Fig. 8 shows a downward spike at 2.3 Gyrs. This spike arises because CO2 is photolyzed most strongly for atmospheres with intermediate O2 concentrations resulting in high O2 production rates in the upper atmosphere (whereas O2 destruction is not as strong as before the GOE in the lower atmosphere). Therefore, the net (P − L) chemical change is not as negative as for the other scenarios considered. The altitude dependence of the net (P − L) chemical change of O2 from chemical production and destruction for 1 and 100 PAL CO2 is depicted in Fig. 9 for an Archean atmosphere scenario with 10−6 PAL O2 (tgeo = 2.69 Gyrs ago). In the case of 100 PAL CO2 the production of O2 is enhanced in the upper atmosphere because of CO2 photolysis producing O and, hence, O2 . This enhanced stratospheric production and subsequent diffusion of O2 to lower atmospheric layers results in the increased O2 profile above 0.1 bar shown in Fig. 7. In the lower atmosphere, however, the O2 destruction is increased due to enhanced abundance of NOx as a result of increased CO from photolytic production. CO is a main sink for OH hence the reaction NO2 + OH + M → HNO3 + M is less efficient, yielding an increased


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19

403

concentration of NO2 . Therefore, the O2 concentration slightly decreases towards higher pressures (see blue dotted line in Fig. 7 and Section 7.1.3 for detailed chemical analysis) due to NOx catalyzed O2 destruction. The resulting modulus of the column integrated net chemical change |(P − L)| for an Archean atmosphere with 100 PAL CO2 increases by 2.8% and yields an enhanced atmospheric O2 surface flux as shown in Fig. 8 (dotted line) before the GOE. This suggests that high amounts of CO2 in the atmosphere of early Earth before the GOE counteract the accumulation of O2 in the atmosphere from a given photosynthetic biosphere due to opposing, in-situ chemical processes, i.e. O2 production in the upper atmosphere due to CO2 photolysis but O2 destruction in the lower atmosphere due to catalytic cycles.

404

7.1.3 Pathway Analysis with Respect to O2

394 395 396 397 398 399 400 401 402

405 406 407 408 409 410 411 412 413 414 415 416 417 418

We now present a detailed chemical pathway analysis of O2 production and destruction in the atmosphere of early Earth with 10−6 PAL O2 and 100 PAL CO2 . Here, we quantify the dominant production and destruction pathways of O2 by applying PAP (for details see Section 4) to early Earth with a surface O2 vmr of 10−6 PAL O2 (tgeo = 2.69 Gyrs) and 100 PAL CO2 . Our study represents the first application of PAP in the context of early Earth with low O2 and high CO2 . As in paper I, we consider only pathways with an individual contribution larger than 1.5% to the total column production (or destruction) of O2 . We use the same pathway numbering as in paper I. All new pathways are indicated in the tables with a star (?). Furthermore, CO was excluded as a branching point species in order to ensure a consistent treatment of all O2 production and destruction pathways at every atmospheric height. This means that partial pathways producing CO are not connected with partial pathways consuming CO. Hence, CO is regarded as a potential source and sink for CO2 in every atmospheric layer.



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20 Tab. 2: Summary of dominant chemical production (P) pathways of O2 found by PAP and their percentage contribution (PC) to the total column-integrated O2 production rate for an Archean atmosphere with a surface O2 vmr of 10−6 PAL O2 and surface CO2 vmr of 100 PAL CO2 . Only pathways that contribute > 1.5% of the total column-integrated production rate are shown. In addition we show the absolute change (AC) in % (difference P C2 − P C, where P C2 denote values for the 1 PAL CO2 scenario given in Tab. 9 (column 6)). For an explanation of classes and subclasses see Appendix. class subclass number PA PA1 P2 P8?

pathway

2·(CO2 + hν → CO + O) O + O 2 + M → O3 + M OH + O → H + O2 H + O3 → OH + O2 net: 2CO2 → O2 + 2CO

PA2 P4 P3 PA3 P9?

PB

2·(CO2 + hν → CO + O) NO2 + O → NO + O2 NO + O + M→ NO2 + M net: 2CO2 → O2 + 2CO

PB1 P6 PB2 P5

419 420 421 422 423 424 425

PC [%] 27.3 21.8 5.4

AC [%] -32.3 -4.2

45.3 42.7 2.6 11.3 11.3

20.0 38.1 -18.1

2.1 2.1 11.0 11.0

-2.6 -2.6 5.3 5.3

Production pathways and their altitude dependence The major column-integrated chemical O2 production pathways found by PAP and their percentage contribution to the total column-integrated O2 production rate are summarized in Tab. 2 for an early Earth atmosphere with 10−6 PAL O2 and 100 PAL CO2 (see Tab. 9 for comparison with 1 PAL CO2 ). For 100 PAL CO2 PAP finds a new subclass PA3 which is catalyzed by NOx because of enhanced NOx compared to 1 PAL CO2 . This is due to enhanced CO


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21

Fig. 10: Selected OH profiles of early Earth analog atmosphere scenarios calculated with the fully updated CAB model (CABMRACO2). 426 427 428 429 430 431 432 433 434 435 436 437 438 439 440 441 442 443 444

from CO2 photolysis. CO is the main sink for OH. Fig. 10 shows that for a 100 PAL CO2 atmosphere with 10−6 PAL O2 OH is reduced below 3·10−4 bar. The decrease in OH leads to an increase in NOx via a slowing in the rate of reactions such as e.g. NO2 + OH + M → HNO3 + M (see Fig. 11). New pathways P8? and P9? now appear. Pathways P2, P3, P4, P5 and P6 are also present in Tab. 9 but their relative contributions are very different. Pathways P1 and P7 (in which CO2 photolysis produces O(1 D) followed by subsequent HOx reactions) are now negligible due to less far-UV radiation than for Archean Earth with 10−6 PAL O2 and 1 PAL CO2 . Note that P8? is closely linked to P7 and only differs in O production from the triplet CO2 photodissociation channel. Additionally, P3 which also produces O(1 D) from CO2 using the singlet photodissociation channel is still operating but now proceeds higher in the atmosphere than P1 and P7. In an early Earth atmosphere with 100 PAL CO2 the largest contribution originates from subclass PA2 (45.3%) which is 20% larger compared to the scenario with 1 PAL CO2 . The contribution of pathway P4 is 38.1% larger because of enhanced CO2 resulting in enhanced O production from photolysis whereas the contribution of pathway P3 producing O(1 D) is 18.1% smaller because of less far-UV radiation penetrating the atmosphere. The second largest contribution arises from subclass



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Fig. 11: Selected NO profiles of early Earth analog atmosphere scenarios calculated with the fully updated CAB model (CABMRACO2). 445 446 447 448 449 450 451 452 453 454 455 456 457 458 459 460 461 462 463

PA1 (27.3%) which is 32.3% smaller compared to the scenario with 1 PAL CO2 . Despite higher CO2 concentration the contribution of pathway P2 is 4% smaller due to less OH below 3·10−4 bar (see Fig. 10) catalyzing this pathway. The new pathways P8? is initiated by CO2 photolysis producing O and involves HOx and O3 which catalyze the production of O2 . The third largest contribution arises from subclass PA3 (11.3%) with the new pathway P9? which is initiated by CO2 photolysis producing O2 from subsequent reactions involving NOx species, followed by subclass PB2 (11%) where the rate of pathway P5 is 5.3% larger compared to the 1 PAL CO2 scenario due to the delivery of O from upper layers by diffusion. Fig. 12 shows the altitude dependence of the O2 production pathways for an Archean atmosphere with 10−6 PAL O2 and 100 PAL CO2 . The total production rate of O2 (solid black line) is also plotted. The sum of all pathways with an individual contribution > 1.5% of the total column-integrated O2 production rate found by PAP amounts to 99.2% of the total column-integrated O2 production rate which is enhanced by 261% amounts to 2.5·1011 molec./cm2 s compared to an early Earth atmosphere with 1 PAL CO2 . Fig. 12 indicates that the strongest production of O2 occurs in the upper atmosphere mainly from pathways initiated by CO2 photolysis where UV radiation is strong. A small contribution to the total production of O2 arises in the lower atmosphere from pathway P9? where a maximum in NO arises


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23

Fig. 12: Altitude dependence of O2 production pathways (see Tab. 2) for an early Earth atmosphere with a surface O2 vmr of 10−6 PAL O2 and a surface CO2 vmr 100 PAL CO2 . Additionally, the total production rate of O2 (solid black line) is plotted. 464 465 466 467 468

(see Fig. 11). The new pathway P8? which uses the triplet CO2 photodissociation channel to produce O, occurs mainly in the upper atmosphere where far-UV radiation is suppressed stronger than UVC radiation and, hence, becomes more dominant relative to P7 which contains the singlet CO2 photodissociation channel producing O(1 D). The same applies for P1 relative to P2 and P3 relative to P4.

Tab. 3: As for Tab. 2 but for O2 destruction (L) classes and pathways. For an explanation of classes and subclasses see Appendix. class subclass number LA LA1 L6 L12?

pathway

2·(H + O2 + M → HO2 + M) 2·(NO + HO2 → NO2 + OH) NO2 + hν → NO + O 2·(CO + OH → CO2 + H) NO2 + O → NO + O2 net: O2 + 2CO → 2CO2


PC [%] 63.5 41.4 11.5

AC [%] 18.4 36.0


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24 class subclass number L1 L2 LA2 L3 LB LB1 L13?

pathway

3·(NO2 + hν → NO + O ) 3·(HO2 + O → OH + O2 ) 2·(CH4 + OH → CH3 + H2 O) 2·(CH3 + O2 + M → CH3 O2 + M) 2·(CH3 O2 + NO → H3 CO + NO2 ) 2·(H3 CO + O2 → H2 CO + HO2 ) 2·(H2 CO + OH → H2 O + HCO) 2·(HCO + O2 → HO2 + CO) NO + HO2 → NO2 + OH) net: 3O2 + 2CH4 → 4H2 O + 2CO

LB2 ?

L14

NO2 + hν → NO + O HO2 + O → OH + O2 CH4 + OH → CH3 + H2 O CH3 + O2 + M → CH3 O2 + M CH3 O2 + NO → H3 CO + NO2 H3 CO + O2 → H2 CO + HO2 H2 CO + hν → H2 + CO net: O2 + CH4 → H2 O + H2 + CO

L15?

NO2 + hν → NO + O NO + HO2 → NO2 + OH HO2 + O → OH + O2 2·(H2 + OH → H2 O + H) 2·(H + O2 + M → HO2 + M) net: O2 + 2H2 → 2H2 O

LC

469 470 471 472 473 474

PC [%] 7.5 3.1 4.7 4.7 3.7 3.7

AC [%] -20.1 -5.8 -8.2 -3.1 -5.4

1.7 1.7

-0.8

1.9 1.9

-1.2

Destruction pathways and their altitude dependence The major column-integrated chemical O2 destruction pathways found by PAP and their percentage contribution to the total colum-integrated O2 destruction rate are summarized in Tab. 3 for an early Earth atmosphere with 10−6 PAL O2 and 100 PAL CO2 (see Tab. 10 for comparison to 1 PAL CO2 ). The total column-integrated destruction rate of O2 increases by 64.3% and amounts to 2.6·1011 molec./cm2 s compared to the 1 PAL


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25

Fig. 13: Same notation as for Fig. 12 but for destruction (see also Tab. 3). Note, the destruction of O2 is composed of a large number of pathways individually contributing less than 1.5% (white area). 475 476 477 478 479 480 481 482 483 484 485 486 487 488 489 490 491 492 493 494 495

CO2 scenario. Pathways L1, L2, L3 and L6 are also present in an early-Earth analog atmosphere around the Sun presented in Tab. 10 although their relative contributions are very different. Four new pathways, namely L12? , L13? , L14? and L15? arise. In the Archean atmosphere with 10−6 PAL O2 and 100 PAL CO2 the largest contribution arises from subclass LA1 (63.5%) which is 18.4% higher compared to the scenario with 1 PAL CO2 . Within LA1, pathway L6 dominates the columnintegrated O2 destruction with 41.4% (36% higher compared to the scenario with 1 PAL CO2 ) because of enhanced NOx species whereas this pathway is rather minor for the corresponding scenario with 1 PAL CO2 . The contributions of pathways L1 and L2 are both decreased because of less OH in the lower atmosphere. The new pathway L12? with a relative contribution of 11.5% arises due to the higher abundance of tropospheric NOx . For LA2 the largest contribution arises from L3 (4.7%) which is 3% smaller in comparison to the 1 PAL CO2 scenario because of the increase in total O2 destruction rate and less OH. The new pathways L13? , L14? and L15? are initialized by the photolysis of NO2 and catalyzed by HOx and NOx . The altitude dependence of the presented destruction pathways for an Archean atmosphere with 10−6 PAL O2 and 100 PAL CO2 is given in Fig. 13. The total destruction rate of O2 is also shown (solid black line). The sum of all pathways with an individual contribution > 1.5% of the total column-integrated O2 destruction rate found by PAP amounts to 79.7% of the total column-integrated O2 destruction rate. The strongest destruction of O2 occurs in the lower atmosphere where destruction is



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26

Fig. 14: Calculated relative change in % of the net primary productivity from oxygenic photosynthesis, N , for early Earth atmosphere scenarios with 10 PAL CO2 (dashed line) and 100 PAL CO2 (dotted line) compared with an atmosphere having 1 PAL CO2 as a function of geological time tgeo .

507

composed of a large number of pathways individually contributing less than 1.5% (white area in Fig. 13) followed by a smaller contribution in the upper atmosphere. In summary, for an Archean atmosphere with 10−6 PAL O2 and 100 PAL CO2 , O2 is produced on the one hand more strongly in-situ from CO2 photolysis followed by catalytic reactions with HOx , NOx and O3 in the upper atmosphere. However, on the other hand CO produced from CO2 photolysis participates in HOx and NOx catalyzed pathways which destroy O2 mainly in the lower atmosphere. The contributions of CH4 and H2 oxidation are rather small. The increase in column-integrated production is more than compensated by the increase in column-integrated destruction and, hence, the absolute (negative) value of the net (P − L) chemical change for an early Earth atmosphere with 100 PAL CO2 increased by 2.8% amounting to a difference of about 9·109 molec./cm2 s relative to the scenario with 1 PAL CO2 .

508

7.2 Biogeochemical Modeling

496 497 498 499 500 501 502 503 504 505 506

509 510 511 512 513 514 515 516 517 518

Fig. 14 presents the relative change of the input from oxygenic photosynthesis, N , in % over geological time for early Earth atmospheres with 10 and 100 PAL CO2 compared with an atmosphere having 1 PAL CO2 . Before the GOE where N is sensitive to the O2 surface flux and the chemical nature of the atmosphere (see paper I), values are slightly increased for an atmosphere with 100 PAL CO2 compared to an early Earth atmosphere with 1 PAL CO2 . The increase for the 10 PAL CO2 scenario is rather small. A drop in N occurs around the GOE because the increase in O2 surface flux is less strong than for the corresponding scenario with 1 PAL CO2 (see Fig. 8). The 100 PAL CO2 scenario calculates a sharp drop in the O2 surface flux (see discussion on O2 above for de-


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27

Fig. 15: Selected p − T profiles of early Earth analog atmosphere scenarios calculated with the fully updated CAB model (CABMRACO2).

527

tails). After the GOE, N increases with increasing O2 content (i.e. going forward in time) because of increasing surface temperatures and pressure from evaporating H2 O. Therefore, high CO2 amounts in our scenarios before the GOE could counteract the accumulation of O2 in the atmosphere from a given photosynthetic biosphere due to the more destructive nature of the atmosphere towards O2 despite higher abiotic production from CO2 photolysis. This destructive nature is driven by enhanced CO removing O2 into CO2 catalyzed by higher tropospheric NOx due to decreased OH concentrations below 3·10−4 bar.

528

8

529

8.1

530

8.1.1 Climate Responses

519 520 521 522 523 524 525 526

531 532 533

Sensitivity with Respect to Variation of Surface Pressure Atmospheric Modeling

Fig. 15 depicts the impact of changing the background surface pressure, p0 , in the range implied by new proxy data upon p − T profiles for selected early Earth atmosphere scenarios with 1 PAL CO2 before, during and after the GOE. Results



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Fig. 16: Surface temperature evolution of early Earth analog atmosphere scenarios calculated with the fully updated CAB model (CABMRACO2). Additionally indicated as a horizontal line is the freezing temperature of H2 O of T = 273.15 K.

547

suggest that the impact is relatively modest. As a first step, we chose to perform the surface pressure variation for atmospheres with 1 PAL CO2 in order to minimize the number of parameters varied at a time, although scenarios with larger amounts of CO2 could be considered for future work. As expected surface temperatures (see Fig. 16) increase with surface pressure due to increased column mass and, hence, enhanced greenhouse warming. Recent constraints 53,55,56 (derived by three different methods) suggest psurf ≤ 0.5 bar on Archean Earth. Such a low surface pressure suggests weakened greenhouse heating which aggravates the FYSP and suggests that more greenhouse heating would be needed to maintain surface temperatures above freezing. Note again that in the case of the scenarios prior to 1.3, 1.9 and 2.2 Gyrs ago with 0.5, 1 and 1.5 bar surface pressure respectively, global surface temperatures are below 0◦ C implying no habitable surface conditions. However, such conditions could still exist locally as previously discussed (see Section 7).

548

8.1.2 Photochemical Responses

534 535 536 537 538 539 540 541 542 543 544 545 546

549 550 551 552 553 554 555 556 557

Fig. 17 shows the UVA, UVB and UVC radiation profiles for the same scenarios as in Fig. 15. In all scenarios increasing the surface pressure (and, hence, column mass) leads to an increase in UVA radiation (upper panel) at the TOA due to enhanced scattering from below but a decrease at the surface due to enhanced absorption. In the case of UVB and UVC, no strong deviations occur for scenarios during and after the GOE (red and black lines in middle and lower panel of Fig. 17) whereas before the GOE (blue lines in Fig. 17) the same response as for the UVA radiation was calculated. This can be attributed to increased amount of O3 with increasing pressure in the atmospheres before the GOE. The influence of vary-


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Fig. 17: UVA (top), UVB (middle) and UVC (bottom) radiation profiles in W/m2 of early Earth analog atmospheric scenarios calculated with the fully updated CAB model (CABMRACO2).



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Fig. 18: Ratio (surface/TOA) radiation fluxes, RUV , for UVA (black), UVB (red) and UVC (green) of early Earth analog atmosphere scenarios calculated with the fully updated CAB model (CABMRACO2) as a function of geological time tgeo . 558 559 560 561

562 563 564 565 566 567 568 569 570 571 572 573 574 575 576 577 578 579 580

ing surface pressure upon the evolution of RU V is shown in Fig. 18. The higher the atmospheric column mass the more efficient radiation is blocked. This effect is strongest for UVA over all timescales considered. For UVB and UVC this effect is strongest before the GOE. Ozone - O3 Fig. 19 presents selected O3 profiles of early Earth atmospheric scenarios before, during and after the GOE. The impact of increasing the surface pressure is strongest before the GOE (blue lines) where O3 differs by one order of magnitude for p > 10−2 bar. This increase in O3 in the lower atmosphere was due to enhanced smog production. This in turn was related to an increase in UV radiation (see Fig. 17) producing more NOx from its reservoir species hence stimulating the smog-mechanism. The UV response was discussed above. Enhanced UV can lead to release of NOx from reservoirs such as HNO3 . The same effect is observed for scenarios during and to some extent after the GOE although changes occur mainly for p > 0.1 bar. The evolution of the O3 column amount for different surface pressures is given in Fig. 20. In order to visualize the impact of increasing surface pressure before and after the GOE we show here two panels having different y-axes, i.e. linear (upper) and logarithmic (lower). Largest differences occur before the GOE where for example the O3 column is decreased by 85% for the scenario with 0.5 bar compared to 1 bar at 2.7 Gyrs. After the GOE, reducing the surface pressure has only a minor effect (i.e. -10.9% for 0.5 bar and 3.3% for 1.5 bar for 1 PAL O2 ) upon the O3 layer. This is consistent with established feedbacks 76 in which O3 column values are preserved over a range of O2 amounts.


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31

Fig. 19: Selected O3 profiles of early Earth analog atmosphere scenarios calculated with the fully updated CAB model (CABMRACO2). 581 582 583 584 585 586 587 588 589 590 591 592 593 594 595 596 597 598 599

Molecular oxygen - O2 Fig. 21 shows selected O2 profiles for the early Earth scenarios with different surface pressures. O2 is decreased by up to one order of magnitude in the case of the Archean scenario with a surface pressure of 0.5 bar. For the other scenarios changes are negligible. To understand the Archean decrease at 0.5 bar (and the slight increase in the case of the Archean atmosphere with 1.5 bar surface pressure) we apply PAP for a detailed chemical analysis of the O2 sources and sink pathways (see below). The evolution of the atmospheric O2 surface flux which is needed to maintain a specific O2 surface mixing ratio for early Earth atmospheres with 0.5, 1 and 1.5 bar surface pressure is shown in Fig. 22. Before the GOE, the O2 surface flux for the 0.5 (1.5) bar atmosphere is slightly increased (decreased) by about 1%. This indicates that the net (P − L) chemical change is more (less) negative and, hence, these atmospheres are chemically somewhat more (less) destructive to O2 (see below). This trend is however reversed shortly before the GOE i.e. low pressure atmospheres are less destructive towards O2 . Fig. 22 shows that the proposed low surface pressure of 0.5 bar for early Earth before the GOE at 2.7 Gyrs ago would have only a negligible effect on the NPP of a given photosynthetic biosphere. After the GOE, the effect of changing pressure is more noticeable. Here, a lowering of the surface pressure leads to a lowering in O2 surface flux and, hence, less net chemical



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Fig. 20: Evolution of O3 column [DU] amount of early Earth analog atmosphere scenarios calculated with the fully updated CAB model (CABMRACO2). Top: linear y-axis. Bottom: logarithmic y-axis.


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33

Fig. 21: Selected O2 profiles of early Earth analog atmosphere scenarios calculated with the fully updated CAB model (CABMRACO2).

Fig. 22: Evolution of O2 surface flux [Tg/yr] of early Earth analog atmosphere scenarios calculated with the fully updated CAB model (CABMRACO2).



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Fig. 23: O2 net chemical change (P − L) for early Earth analog atmosphere scenarios with a surface O2 vmr of 10−6 PAL O2 calculated with the fully updated CAB model (CABMRACO2) for surface pressures of 0.5 bar (dashed), 1 bar (solid) and 1.5 bar (dotted). The vertical dot-dashed line indicates the zero line.

613

O2 destruction and, hence, the NPP would be reduced. The altitude dependence of the net (P − L) chemical change of O2 from chemical production and destruction for Archean atmospheres with 10−6 PAL O2 (tgeo = 2.69 Gyrs ago) and 0.5, 1 and 1.5 bar surface pressure is depicted in Fig. 23. The main differences occur for p > 0.01 bar. At a given pressure level, O2 is destroyed more strongly in the scenario with 0.5 bar surface pressure compared to the scenario with 1 bar, resulting in a decreased O2 profile (see Fig. 21). For the atmosphere with 1.5 bar surface pressure an O2 production region arises between 0.3 to 0.1 bar. This arose mainly due to increased NOx and enhanced O from CO2 photolysis (see analysis below). We now present a detailed chemical analysis of the O2 production and destruction pathways in the atmosphere of early Earth with 10−6 PAL O2 and 0.5 and 1.5 bar surface pressure. Here, we quantify the dominant O2 production and destruction pathways by applying PAP.

614

8.1.3 Pathway Analysis with Respect to O2

600 601 602 603 604 605 606 607 608 609 610 611 612

615 616 617 618

PAP (for details see Section 4) is applied to early Earth with a surface O2 vmr of 10−6 PAL O2 (tgeo = 2.69 Gyrs), 1 PAL CO2 and a surface pressure of 0.5 and 1.5 bar. Our study represents the first application of PAP in the context of early Earth with low O2 and varying surface pressure.


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35 Tab. 4: Summary of dominant chemical production (P) pathways of O2 found by PAP and their percentage contribution to the total column-integrated O2 production rate for an Archean atmosphere with a surface O2 vmr of 10−6 PAL O2 , a surface CO2 vmr of 1 PAL CO2 and 0.5 or 1.5 bar surface pressure. Only pathways that contribute > 1.5% of the total column-integrated production rate are shown. In addition we show the absolute change (AC) in % (difference P C2 − P C, where P C2 denote values for the 1 bar surface pressure scenario given in Tab. 9 (column 6)). For further explanation of classes and subclasses see Appendix. 0.5 bar 0.5 bar class subclass number PC [%] AC [%] PA PA1 59.9 0.3 P1 33.2 1.8 P2 24.1 -1.9 P7 2.6 0.3 PA2 24.3 -1.0 P3 22.0 1.3 P4 2.3 -2.3 PA3 P9? PB PB1 5.0 0.3 P6 5.0 0.3 PB2 6.0 0.3 P5 6.0 0.3 619 620 621 622 623 624 625 626 627 628 629

1.5 bar 1.5 bar PC [%] AC [%] 58.3 -1.3 28.5 -2.9 27.8 1.8 2.0 -0.3 26.2 0.9 18.6 -2.1 7.6 3.0 2.1 2.1 4.3 -0.4 4.3 -0.4 5.3 -0.4 5.3 -0.4

Production pathways and their altitude dependence The major column-integrated chemical O2 production pathways found by PAP and their percentage contribution to the total column-integrated O2 production rate for an early Earth atmosphere with 10−6 PAL O2 , 1 PAL CO2 are summarized in Tab. 4 for p0 = 0.5 bar and p0 = 1.5 bar (for comparison see Tab. 9 which shows results for 1 bar). Our study suggests that the same pathways operate as for an early Earth atmosphere with 1 bar surface pressure but with slightly different contributions to the total column-integrated O2 production rates. In the case of the 1.5 bar scenario we find a new pathway, namely P9? , which can be attributed to subclass PA3 where the net reaction 2CO2 → O2 + 2CO is initiated by CO2 photolysis and catalyzed by NOx (see Tab. 2). On increasing the surface pressure from 0.5 to 1.5 bar, the largest individual difference



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36

Fig. 24: Altitude dependence of O2 production pathways (see Tab. 4) for an early Earth atmosphere with a surface O2 vmr of 10−6 PAL O2 , a surface CO2 vmr of 1 PAL CO2 and a surface pressure of 0.5 bar (top) and 1.5 bar (bottom). The solid black line shows the total production rate of O2 .


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37

645

occurs for pathway P4 where the contribution increases by 5.3% due to the increase in UV radiation (on a given pressure levels) throughout the atmosphere (see Fig. 17) producing O from CO2 photolysis. Fig. 24 shows the altitude dependence of the O2 production pathways for Archean atmospheres having 10−6 PAL O2 and 1 PAL CO2 but with different surface pressures (top: 0.5 bar, bottom: 1.5 bar). The total O2 production rates (solid black lines) are additionally plotted. Compared with the corresponding 1 bar surface pressure scenario this value is decreased (increased) by 5.9 (10.6)% for the scenario with 0.5 (1.5) bar surface pressure. The sum of all pathways with an individual contribution > 1.5% of the total column-integrated O2 production rate found by PAP amounts to 98.9% and 98.6% of the total column-integrated O2 production rate for the scenarios with 0.5 and 1.5 bar, respectively. Overall, the two panels are broadly similar which suggests that the effect of changing p0 upon the O2 production pathways is only modest. Fig. 24 indicates that the largest increase in O2 production with surface pressure occurs in the middle atmosphere mainly arising from pathways P2, P4 and P9? .

646

Destruction pathways and their altitude dependence

630 631 632 633 634 635 636 637 638 639 640 641 642 643 644


1 2 3 4 5 6 7 8 9 10 11 12 13 14 15 16 17 18 19 20 21 22 23 24 25 26 27 28 29 30 31 32 33 34 35 36 37 38 39 40 41 42 43 44 45 46 47 48 49 50 51 52 53 54 55 56 57 58 59 60

class subclass number LA1 LA L1 L2 L8 L6 LA2 L3 L5 LB LB1 L7 L4 L16? 2·(H2 O + hν → H + OH) 2·(H + O2 + M → HO2 + M) 3·(2HO2 → H2 O2 + O2 ) 3·(H2 O2 + hν → 2OH) 2·(CH4 + OH → CH3 + H2 O) 2·(CH3 + O2 + M → CH3 O2 + M) 2·(CH3 O2 + OH → H3 CO + HO2 ) 2·(H3 CO + OH → H2 CO + H2 O) 2·(H2 CO + OH → H2 O + HCO)

pathway

0.5 bar 0.5 bar 1.5 bar 1.5 bar PC [%] AC [%] PC [%] AC [%] 47.7 2.6 44.9 -0.2 26.1 -1.5 27.2 -0.4 10.0 1.2 7.6 -1.3 8.1 4.8 2.8 -0.5 3.4 -2.0 7.2 1.8 10.6 -2.3 13.4 0.5 5.7 -2.1 8.3 0.5 4.9 -0.3 5.1 -0.1 11.5 2.5 8.8 -0.3 5.6 1.7 2.4 -1.5 3.8 -1.3 4.8 -0.3 2.1

Tab. 5: As for Tab. 4 but for O2 destruction (L) classes and pathways. For an explanation of classes and subclasses see Appendix.



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38


LD

LC

LB2

L11

L9

L10

L17?

class subclass number

pathway 2·(HCO + O2 → HO2 + CO) net:3O2 + 2CH4 → 4H2 O + 2CO 2·(H + O2 + M → HO2 + M) 2HO2 → H2 O2 + O2 H2 O2 + hν → 2OH 2·(CH4 + OH → CH3 + H2 O) 2·(CH3 + O2 + M → CH3 O2 + M) 2·(CH3 O2 + HO2 → CH3 OOH + O2 ) 2·(CH3 OOH + hv → H3 CO + OH) 2·(H3 CO + O2 → H2 CO + HO2 ) 2·(H2 CO + hv → HCO + H) 2·(HCO + O2 → HO2 + CO) 2·(OH + HO2 → H2 O + O2 ) net:3O2 + 2CH4 → 4H2 O + 2CO

1 2 3 4 5 6 7 8 9 10 11 12 13 14 15 16 17 18 19 20 21 22 23 24 25 26 27 28 29 30 31 32 33 34 35 36 37 38 39 40 41 42 43 44 45 46 47 48 49 50 51 52 53 54 55 56 57 58 59 60

3.2 3.2 1.8 1.8

0.1 0.1 -0.2 -0.2

3.1 3.1 2.7 2.7 2.0 2.0

1.6

0.7 0.7 -0.4 -0.4 0.0 0.0

0.5 bar 0.5 bar 1.5 bar 1.5 bar PC [%] AC [%] PC [%] AC [%]

Page 39 of 71 ACS Earth and Space Chemistry

39


Page 40 of 71

40

Fig. 25: Selected OH profiles of early Earth analog atmosphere scenarios calculated with the fully updated CAB model (CABMRACO2). 647 648 649 650 651 652 653 654 655 656 657 658 659 660 661 662 663 664 665

Tab. 5 summarizes the major column-integrated chemical O2 destruction pathways found by PAP and their percentage contribution to the total column-integrated O2 destruction rate for an early Earth atmosphere with 10−6 PAL O2 , 1 PAL CO2 for p0 = 0.5 bar and p0 = 1.5 bar (see Tab. 10 for comparison for 1 bar). We find the same pathways as for the 1 bar surface pressure scenario but with slightly different contributions. In addition, new CH4 oxidation pathways belonging to subclass LB1, namely L16? and L17? , were identified. On increasing the surface pressure from 0.5 to 1.5 bar, the largest individual difference occurs for pathway L8 where the contribution decreases by 5.3% because of decreasing H, HOx and H2 O concentrations below 0.01 bar (see for e.g. OH Fig. 25). Fig. 26 shows the altitude dependence of the destruction pathways for the Archean atmosphere with 10−6 PAL O2 and 1 PAL CO2 at different surface pressures (top: 0.5 bar, bottom: 1.5 bar). As for the production case (see Fig. 24), the upper and lower panels in Fig. 33 are similar which suggests that varying p0 has only a modest effect on the chemical destruction pathways for atmospheric O2 . The total destruction rate of O2 is also shown (solid black lines). The total columnintegrated destruction rate of O2 is decreased (increased) by 0.7 (1.8)% for the 0.5 (1.5) bar scenario compared to an early Earth atmosphere with a surface pressure of 1 bar. In the case of the scenario with 0.5 (1.5) bar surface pressure the sum of


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41

Fig. 26: Same notation as for Fig. 12 but for destruction (see also Tab. 5). Note, the destruction of O2 is composed of a large number of pathways which individually contribute less than 1.5% to the total O2 destruction rate (white area).



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42

Fig. 27: Calculated relative change in % of the input from oxygenic photosynthesis, N , for early Earth atmosphere scenarios with 0.5 dashed line and 1.5 bar surface pressure dotted line relative to an atmosphere with psurf = 1 bar as a function of geological time tgeo .

684

all pathways with an individual contribution > 1.5% of the total column-integrated O2 destruction rate found by PAP amounts to 84.7 (84.0)% of the total columnintegrated O2 destruction rate. The strongest destruction of O2 occurs in the lower atmosphere where destruction is caused by a large number of pathways individually contributing less than 1.5%. For a low surface pressure atmosphere the total O2 destruction rate at the surface is stronger due to higher UV radiation whereas for the high surface pressure scenario UV radiation is more absorbed due to more atmospheric mass. In the upper atmosphere (at identical pressure levels) this behavior is reversed (see Fig. 17). In summary, for an Archean atmosphere with 10−6 PAL O2 and 1 PAL CO2 , with increasing surface pressure O2 is produced more strongly in the middle atmosphere (at a pressure level of about 0.3 bar) by either in-situ CO2 photolysis alone, or coupled with catalytic reactions involving HOx . For high surface pressure atmospheres NOx becomes important in catalyzing pathway P9? . The destruction of O2 is dominated by many different oxidation pathways. The decrease (increase) in column-integrated production is mostly canceled out by the decrease (increase) in column-integrated destruction for a 0.5 (1.5) bar surface pressure atmosphere and, hence, the net (P − L) chemical change is similar (< 0.9%) to the early Earth atmosphere with 1 bar surface pressure.

685

8.2

666 667 668 669 670 671 672 673 674 675 676 677 678 679 680 681 682 683

686 687 688

Biogeochemical Modeling

Fig. 27 presents the relative change in % in the input from NPP (N ) over geological time for early Earth atmospheres with psurf = 0.5 bar and psurf = 1.5 bar against an atmosphere with 1 bar surface pressure and 1 PAL CO2 .


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43

Fig. 28: Calculated relative change in % of the input from oxygenic photosynthesis, N , for early Earth with Tocean = 70◦ C compared with Tocean = 25◦ C as a function of geological time tgeo .

701

Before the GOE only a small impact on N is calculated since the change in in-situ chemical net destruction is small in this low O2 regime. Shortly before the GOE, on going forward in time, N starts to increase (decrease) for early Earth with 1.5 (0.5) bar in comparison to scenarios with 1 bar surface pressure because in the intermediate O2 regime the sensitivity of the NPP upon atmospheric chemistry becomes smaller whereas the impact of increasing O2 partial pressure becomes larger. During the GOE the rise (decline) of N is enhanced due to the increasing O2 partial pressure and converges towards a 120% increase (75% decrease) until present. The increase (decrease) is additionally related to warmer (cooler) surface temperatures resulting in more (less) evaporation of H2 O and, hence, surface pressure. The decreasing N for low surface pressure atmospheres before and during the GOE implies that the accumulation of atmospheric O2 up to a given abundance would be achievable for a lower level of photosynthetic activity.

702

9

689 690 691 692 693 694 695 696 697 698 699 700

703 704 705 706 707 708

Sensitivity with Respect to Variation of Ocean Temperature

In this section we analyze the impact of a high ocean temperature on N which have been proposed for early Earth by several geochemical data studies (see Section 1). A first approach we crudely assume no feedback between a hot surface ocean and the atmosphere since parameterizing this would be beyond the scope of our work. Hence, we use the same atmospheric temperature, chemical species profiles and atmospheric O2 surface fluxes as calculated in Section 7 for 1 PAL CO2 .



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44

709

9.1

Biogeochemical Modeling

717

The resulting relative change for N while changing the ocean surface temperature from 25 to the maximum proposed temperature of 70◦ C is depicted in Fig. 28. The strongest effect is observed before the GOE where N is decreased by up to 20%. This is because the O2 solubility is decreased for a higher ocean temperature and, hence, O2 could accumulate in the atmosphere more easily. After the GOE N is no longer sensitive to the atmospheric flux of O2 but rather the O2 partial pressure (see paper I). Changing the solubility constant in this regime therefore has a much weaker effect on the high N values associated with the established biosphere.

718

10

710 711 712 713 714 715 716

719 720 721 722 723 724 725 726 727 728 729 730 731 732 733 734 735 736 737 738 739 740 741 742 743 744

Summary and Conclusion

We have applied the updated CAB model 68 to the early Earth in order to analyze the effect of geologically-constrained environmental variables such as CO2 content, surface pressure and ocean temperature on the atmosphere and most importantly on the corresponding productivity of a photosynthesizing biosphere. The calculated increase of O2 for pressures below 0.3 bar with increasing atmospheric CO2 content (see Fig. 7) is related to the enhanced production from CO2 photolysis in the upper atmosphere. However, O2 destruction in the lower atmosphere is also increased due to enhanced abundance of NOx (see Fig. 11) and CO yielding a slightly decreasing O2 concentration towards higher pressures. The resulting column integrated net (P − L) chemical change for high CO2 atmospheres is increased by 2.8%. Therefore, atmospheres with low surface O2 vmr such as before the GOE but with high amounts of CO2 as proposed by recent studies e.g. 100 PAL CO2 could counteract the accumulation of atmospheric O2 from a given photosynthetic biosphere because of enhanced atmospheric O2 destruction due to these opposing processes. Reducing the surface pressure to 0.5 bar in an Early Earth atmosphere as proposed by several studies in the recent years yields a decreased O2 concentration profile between 0.5 to 0.005 bar (see blue dashed line in Fig. 21) compared to an early Earth atmosphere with higher surface pressures. This is a result of enhanced destruction by enhanced HOx (see Fig. 25 and upper panel in Fig. 26) catalyzing these pathways and lower production (see Fig. 24) due to less UV radiation destroying CO2 in this pressure regime. Lowering the surface pressure to 0.5 bar has a negligible effect on the NPP at 2.7 Gyrs ago but shortly before the Great Oxidation Event NPP is decreased by about 20% (see Fig. 27). This suggests that the accumulation of atmospheric O2 up to a given abundance would be achievable for a lower level of photosynthetic activity.


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45

745 746 747 748 749 750 751 752 753 754 755 756 757 758 759 760 761 762 763 764 765 766 767 768 769 770

We have applied the Pathway Analysis Program to quantify production and destruction pathways of O2 for Archean Earth to analyze the behavior of the O2 profile during different geological constraints. For high CO2 atmospheres we identified additional production pathways which are catalyzed by O3 in the upper atmosphere and in the lower atmosphere by NOx (see Tab. 7 and Fig. 17). For the destruction of O2 we find four new pathways due to the higher abundance of NOx in the lower atmosphere (see Tab. 8 and Fig. 20). With increasing surface pressure the production of O2 is enhanced and a NOx catalyzed pathway appears in the lower atmosphere (see Tab. 9 and Fig. 32). In the case of the destruction of O2 new complex CH4 and H2 oxidation pathways arise in the lower atmosphere for the low and high surface pressure atmospheres, although their contributions are rather minor (see Tab. 10 and Fig. 33). If we assume that the ocean was hot as proposed by some recent studies then the NPP from oxygenic photosynthesis is reduced by about 25% due to lower O2 solubility. This enabled O2 to accumulate in the atmosphere more easily. In summary, it can be stated that, whereas uncertainties in pressure and ocean temperature have only a modest effect on early Earth’s climate and composition, uncertainties in CO2 have a large effect, especially in the stratosphere. This has repercussions for atmospheric escape and planetary evolution on early Earth and Earth-like planets. A new generation of atmospheric models (such as that used in this study) with coupled biogeochemistry and pathway diagnostics are needed to understand the complex and subtle interactions which are at work. The photochemistry analysis has revealed that whereas O2 is mainly produced in the middle-upper atmosphere, it is destroyed mainly in the lower atmosphere. The analysis has also revealed the key role played by HOx and NOx and an important implication is that the photochemical responses and reservoir species need further investigations.



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46

771 772 773 774 775

Acknowledgments: Stefanie Gebauer acknowledges support by the DFG project GZ:GR2004/2-1 of the SPP 1833 ”Building a Habitable Earth”. Author Disclosure Statement: No competing financial interests exist.


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47

776

A

CAB Model Validation and Comparison with Previous Model Version

777

778

A.1

Modern Earth

779

A.1.1 Atmosphere Module Tab. 6: Biogenic surface fluxes of CH4 , CO, N2 O and CH3 Cl for modern Earth calculated with the CAB model of paper I (CABRRTM), with MRAC-complete without the new O2 photolysis scheme (CABMRAC) and the fully updated CAB model (CABMRACO2) compared to observations 91,92 and references therein. Values in brackets represent the natural (without anthropogenic) component only. Fluxes are given in Teragrams (Tg (1012 g))/yr. species CABRRTM CABMRAC CABMRACO2 Observation species [Tg/yr] [Tg/yr] [Tg/yr] [Tg/yr] CH4 474 482 360 500-600 (150-270) CO 1796 1800 1476 1060 N2 O 13.5 13.4 14.8 21.1-115 (14.5-52.8) CH3 Cl 3.4 3.4 2.7 2.9 (2.7)

780 781 782 783 784 785 786 787 788 789 790 791 792 793 794 795 796

The resulting changes for the biogenic surface fluxes of CH4 , CO, N2 O and CH3 Cl required in order to reproduce the modern Earth are given in Tab. 6. All biogenic surface fluxes are reasonably close to the observed values 91,92 and references therein, however the deviation for CH4 is somewhat larger, whereas the surface flux for CO and CH3 Cl are closer to the observed value than before. Due to less chemical destruction in the atmosphere which is associated with lower hydroxyl radical (OH) concentrations (see Fig. 30), lower biogenic surface fluxes of CH4 , CO, and CH3 Cl are sufficient to maintain the atmospheric abundance of these species. A slightly increased N2 O surface flux is needed to maintain the original N2 O volume mixing ratio (vmr) because of increased photolytic destruction associated with stronger UV due to less O3 in the upper atmosphere. The H2 deposition velocity changed from vdep = 2.63 · 10−3 cm s−1 in paper I to vdep = 2.54 · 10−3 cm s−1 . Fig. 29 shows a good agreement for the temperature profile between CABRRTM (black line) and CABMRACO2 (blue line) except in the uppermost atmospheric layers due to less heating by O3 which is reduced in this pressure range due to less photolytically produced O and excited oxygen [O(1 D)] from O2 and CO2 . Surface UVA and UVB radiation are reduced from about 90 and 2.3 W m−2



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48

Fig. 29: Pressure-temperature (p-T) profiles of modern Earth calculated with the CAB model of paper I (CABRRTM) shown in black, with MRAC-complete without the new O2 photolysis scheme (CABMRAC) shown in red and the fully updated CAB model (CABMRACO2) shown in blue. 797 798 799 800 801 802 803 804 805 806 807

respectively in paper I and CABMRAC to 77 and 1.5 W m−2 in the case of CABMRACO2. This is related to a slightly enhanced O3 profile around the O3 concentration maximum (see Fig. 30 and Tab. 7). UVC radiation at the surface is hardly affected and remains at 0 W m−2 . The modern Earth, global mean observed UVB value for cloud-free conditions of 1.4 W m−2 93 is now better reproduced with the new O2 column dependent S-R cross sections. The impact on biosignatures, related compounds as well as OH is depicted in Fig. 30. Changes are minor for O3 , H2 O, CH4 , N2 O and CH3 Cl. For OH one finds lower concentrations below about 45 km due to a decrease in O(1 D) which weakens the reaction H2 O + O(1 D) → 2OH. Above this height the concentration of OH is enhanced due to e.g. slightly enhanced H2 O. Tab. 7: Column amount [DU] (Dobson unit, 2.69 · 1016 molecules cm−2 ) of selected chemical species for modern Earth calculated with the CAB model of paper I (CABRRTM), with MRAC-complete without the new O2 photolysis scheme (CABMRAC) and the fully updated CAB model (CABMRACO2) species CABRRTM CABMRAC CABMRACO2 O3 305.0 304.8 311.1


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49 species CABRRTM CABMRAC CABMRACO2 H2 O 2.418 · 106 2.441 · 106 2.450 · 106 CH4 1237 1231 1235 N2 O 230.5 230.9 230.1 CH3 Cl 0.358 0.356 0.361

816

Tab. 7 summarizes the changes in column amounts for the selected chemical species O3 , H2 O, CH4 , N2 O and CH3 Cl in Dobson Unit (DU, 2.69 · 1016 molecules cm−2 ). Only minor changes are found. The effect on the atmospheric O2 surface flux which is needed in the model to maintain modern Earth’s surface O2 volume mixing ratio (vmr) decreases from about 3700 Tg O2 /yr as calculated in paper I to 3200 Tg O2 /yr for CABMRACO2 which reflects that the modulus of the column integrated net chemical change, |(P − L)|, is decreased by about 15%. Note that (P − L) increases from a negative value to a less negative value.

817

A.1.2 Biogeochemical Module

808 809 810 811 812 813 814 815

822

The biogeochemistry module of the CAB model calculates the Net Primary Productivity (NPP), N , from oxygenic photosynthesis which is needed to maintain the O2 surface flux calculated by the atmospheric chemistry module for a given surface vmr of O2 . The resulting value for N decreases only by a minor amount (0.005%) compared to paper I.

823

A.2 Archean Earth

824

A.2.1 Atmosphere Module

818 819 820 821

825 826 827 828 829 830 831 832 833 834 835 836

To estimate the impact of the new O2 photolysis scheme in the S-R wavelength range we applied CABMRACO2 to the Archean atmosphere scenario of paper I with a surface vmr of 10−6 PAL O2 . Fig. 31 shows some deviation for the temperature profile between CABRRTM (black line) and CABMRACO2 (dashed blue line) whereas the impact of the new O2 S-R cross sections is mostly negligible (red line). Surface UVA, UVB and UVC radiation are changed from about 74, 12.4 and 2.8 W m−2 (as calculated in paper I and with CABMRAC) to 63, 10.0 and 2.9 W m−2 respectively in the case of CABMRACO2. This is related to an enhanced O3 profile in the middle atmosphere and due to stronger scattering by H2 O which is enhanced in the upper atmosphere compared to CABMRAC and paper I (see Fig. 32).



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50

Fig. 30: Selected species profiles (O3 - upper left, H2 O - upper right, CH4 - middle left, N2 O - middle right, CH3 Cl - lower left and OH - lower right) of modern Earth calculated with the CAB model of paper I (CABRRTM) shown in black, with MRAC-complete without the new O2 photolysis scheme (CABMRAC) shown in red and the fully updated CAB model (CABMRACO2) shown in blue.


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51

Fig. 31: p-T profiles of Archean Earth with 10−6 PAL O2 calculated with the CAB model of paper I (CABRRTM) shown in black, with MRACcomplete without the new O2 photolysis scheme (CABMRAC) shown in red and the fully updated CAB model (CABMRACO2) shown in dashed blue. Note, the red and dashed blue line completely overlap. 837 838 839 840 841 842 843 844 845 846

The rather minor impact on biosignatures and related compounds including OH is depicted in Fig. 32. O3 is slightly decreased in the upper atmosphere due to less photolytic production of O and O(1 D) from O2 and CO2 . In the lower atmosphere O3 is decreased between 5 - 20 km compared to paper I because of enhanced H2 O in this range producing HOx . In the upper atmosphere H2 O, CH4 and CH3 Cl concentrations are higher for CABMRACO2 because of less photolytic destruction also due to the increased zenith angle. Due to less photolytic destruction of O3 in the lower atmosphere, the concentration of O(1 D) and the rate of the N2 O loss reaction N2 O + O(1 D) → 2NO are smaller and, consequently, N2 O is slightly enhanced in this region. Tab. 8: Column amount [DU] (Dobson unit, 2.69 · 1016 molecules cm−2 ) of selected chemical species for Archean Earth with 10−6 PAL O2 calculated with the CAB model of paper I (CABRRTM), with MRAC-complete without the new O2 photolysis scheme (CABMRAC) and the fully updated CAB model (CABMRACO2) species CABRRTM O3 3.169 · 10−2

CABMRAC CABMRACO2 2.355 · 10−2 2.361 · 10−2



Page 52 of 71

52 species H2 O CH4 N2 O CH3 Cl 847 848 849 850 851 852 853 854 855 856 857 858 859

CABRRTM 2.607 · 105 713 0.204 0.121

CABMRAC CABMRACO2 2.835 · 105 2.834 · 105 665 660 0.204 0.285 0.115 0.114

Tab. 8 summarizes the changes in column amounts for the selected chemical species O3 , H2 O, CH4 , N2 O and CH3 Cl in DU. Only minor changes are observed. The effect on the atmospheric O2 surface flux which would be needed to maintain an Archean surface O2 vmr of 10−6 PAL decreases from about 3280 Tg O2 /yr (paper I) to 2640 Tg O2 /yr (this work CABMRACO2) which reflects that the modulus of the column integrated net chemical change (P − L) is reduced by about 21%. This can be attributed to a decrease in the overall O2 production rate, P , by about 18% due to less production of O from CO2 photolysis (see analysis below) and increased OH concentration in the upper atmosphere. The column-integrated O2 destruction rate, L, is reduced by about 19% compared to paper I. This decrease was due to less CO in the lower atmosphere in comparison to paper I (see analysis below and Fig. 33). The decrease in CO arose due to less photolytic production from e.g. H2 CO in the lower atmosphere.


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53

Fig. 32: Selected species profiles (O3 - upper left, H2 O - upper right, CH4 - middle left, N2 O - middle right, CH3 Cl - lower left and OH - lower right) of Archean Earth with 10−6 PAL O2 calculated with the CAB model of paper I (CABRRTM) shown in black, with MRAC-complete without the new O2 photolysis scheme (CABMRAC) shown in red and the fully updated CAB model (CABMRACO2) shown in blue.



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54

Fig. 33: CO profiles of Archean Earth with 10−6 PAL O2 calculated with the CAB model of paper I (CABRRTM) shown in black, with MRAC-complete without the new O2 photolysis scheme (CABMRAC) shown in red and the fully updated CAB model (CABMRACO2) shown in blue.


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55

860

A.2.2 Pathway Analysis of O2 production and destruction Tab. 9: Summary of dominant chemical production (P) classes and pathways of O2 found by PAP and their Percentage Contribution (PC) to the total column-integrated O2 production rate for an Archean atmosphere with a surface O2 vmr of 10−6 PAL O2 calculated by CABRRTM (1) and CABMRACO2 (2). Only pathways that contribute > 1.5% of the total column-integrated production rate are shown. For a clarification of classes and subclasses see Section A.2.2. In addition we show the resulting absolute change (AC), P C2 − P C1, in %. class subclass number PA

PA1 P1

P2

P7

PA2 P3

P4

pathway

PC1 PC2 [%] [%] 58.77 59.61 1 2·(CO2 + hν → CO + O( D)) 28.40 31.37 2·(O(1 D) + N2 → O + N2 ) HO2 + O → OH + O2 OH + O → H + O2 H + O2 + M → HO2 + M net: 2CO2 → O2 + 2CO 2·(CO2 + hν → CO + O) 28.07 25.98 HO2 + O → OH + O2 OH + O → H + O2 H + O2 + M → HO2 + M net: 2CO2 → O2 + 2CO 2·(CO2 + hν → CO + O(1 D)) 2.31 2.26 2·(O(1 D) + N2 → O + N2 ) O + O2 + M → O3 + M OH + O → H + O2 H + O3 → OH + O2 net: 2CO2 → O2 + 2CO 27.47 25.28 1 2·(CO2 + hν → CO + O( D)) 20.65 20.67 2·(O(1 D) + N2 → O + N2 ) O + O + M → O2 + M net: 2CO2 → O2 + 2CO 2·(CO2 + hν → CO + O) 6.82 4.61


AC [%] 0.84 2.97

-2.09

-0.05

-2.19 0.02

-2.21


Page 56 of 71

56 class subclass number

pathway

PC1 [%]

PC2 [%]

AC [%]

3.88 3.88

4.72 4.72

0.84 0.84

5.29 5.29

5.73 5.73

0.44 0.44

O + O + M → O2 + M net: 2CO2 → O2 + 2CO PB

PB1 P6

OH + O → H + O2 H + O2 + M → HO2 + M HO2 + O → OH + O2 net: 2O → O2

PB2 P5

O + O + M → O2 + M net: 2O → O2

Tab. 10: Same notation as for Tab. 9 but for O2 destruction (L) classes and pathways. class subclass number LA

pathway

LA1 L1

L2

L6

L8

2·(H + O2 + M → HO2 + M) HO2 + HO2 → H2 O2 + O2 H2 O2 + hν → OH + OH 2·(CO + OH → CO2 + H) net: O2 + 2CO → 2CO2 2·(H + O2 + M → HO2 + M) HO2 + O → OH + O2 HO2 + hν → OH + O 2·(CO + OH → CO2 + H) net: O2 + 2CO → 2CO2 2·(H + O2 + M → HO2 + M) HO2 + O → OH + O2 NO + HO2 → NO2 + OH 2·(CO + OH → CO2 + H) NO2 + hν → NO + O net: O2 + 2CO → 2CO2 H + O2 + M → HO2 + M H + HO2 → OH + OH 2·(CO + OH → CO2 + H) net: O2 + 2CO → 2CO2


PC1 PC2 AC [%] [%] [%] 45.58 45.07 -0.51 28.47 27.55 -0.92

9.37

8.85

-0.52

4.39

5.37

0.98

3.35

3.30

-0.05

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pathway

LA2 L3

L5

LB

O2 + hν → O + O 2·(HO2 + O → OH + O2 ) 2·(CO + OH → CO2 + H) 2·(H + O2 + M → HO2 + M) net: O2 + 2CO → 2CO2 O2 + hν → O + O(1 D) O(1 D) + N2 → O + N2 2·(HO2 + O → OH + O2 ) 2·(CO + OH → CO2 + H) 2·(H + O2 + M → HO2 + M) net: O2 + 2CO → 2CO2

LB1 L4

L7

H2 O2 + hν → OH + OH 2·(CH4 + OH → CH3 + H2 O) 2·(CH3 + O2 + M → CH3 O2 + M) 2·(CH3 O2 + HO2 → CH3 OOH + O2 ) 2·(CH3 OOH + hν → H3 CO + OH) 2·(H3 CO + O2 → H2 CO + HO2 ) 2·(H2 CO + OH → H2 O + HCO) 2·(HCO + O2 → HO2 + CO) HO2 + HO2 → H2 O2 + O2 net: 3O2 + 2CH4 → 4H2 O + 2CO 2·(H2 CO + OH → H2 O + HCO) 2·(HCO + O2 → HO2 + CO) 3·(HO2 + HO2 → H2 O2 + O2 ) 3·(H2 O2 + hν → OH + OH) 2·(CH4 + OH → CH3 + H2 O) 2·(CH3 + O2 + M → CH3 O2 + M) 2·(CH3 O2 + OH → H3 CO + HO2 ) 2·(H3 CO + O2 → H2 CO + HO2 ) net: 3O2 + 2CH4 → 4H2 O + 2CO

LB2 L10

CH4 + OH → CH3 + H2 O CH3 + O2 + M → CH3 O2 + M CH3 O2 + HO2 → CH3 OOH + O2 CH3 OOH + hν → H3 CO + OH


PC1 PC2 [%] [%] 12.35 12.93 7.37 7.77

AC [%] 0.58 0.40

4.98

5.16

0.18

9.69 5.37

9.05 5.12

-0.64 -0.25

4.32

3.92

-0.4

2.48 2.48

2.45 2.45

-0.03 -0.03


Page 58 of 71


pathway

PC1 [%]

PC2 [%]

AC [%]

2.98 2.98

3.10 3.10

0.12 0.12

1.72 1.72

2.00 2.00

0.28 0.28

H3 CO + O2 → H2 CO + HO2 H2 CO + hν → H2 + CO net: O2 + CH4 → H2 O + H2 + CO LC L9

2·(H + O2 + M → HO2 + M) HO2 + HO2 → H2 O2 + O2 H2 O2 + hν → OH + OH 2·(H2 + OH → H2 O + H) net: O2 + 2H2 → 2H2 O

LD L11

861 862 863 864 865 866 867

868 869 870

CO + OH → CO2 + H H2 O + hν → H + OH 2·(H + O2 + M → HO2 + M) HO2 + HO2 → H2 O2 + O2 net: O2 + H2 O + CO → H2 O2 + CO2

In order to analyze the impact of CABMRACO2 on the O2 production and destruction pathways presented in paper I, we applied PAP to the same Archean atmosphere scenario with a surface vmr of 10−6 PAL O2 . Tabs. 9 and 10 summarize the percentage contribution of the O2 production and destruction classes and individual pathways presented in paper I to the total column-integrated O2 production and destruction rate and the resulting absolute change, respectively. As in paper I, the O2 production pathways can be broadly categorized into two classes, namely: • class PA: O2 is produced by pathways with the net reaction 2CO2 → O2 + 2CO

871

– subclass PA1 (P1, P2, P7): catalyzed by HOx species,

872

– subclass PA2 (P3, P4): without the presence of HOx ,

874

• class PB: O2 is formed by pathways with the net reaction 2O → O2

875

– subclass PB1 (P6): catalyzed by HOx species,

876

– subclass PB2 (P5): without the presence of HOx .

873

877 878

Note that pathways of subclasses PB1 and PA1 and PB2 and PA2 are closely related and only differ in providing O by either in-situ photolysis of CO2 or delivery


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59

Fig. 34: Altitude dependence of O2 production pathways P1 to P7 for an Archean atmosphere with a ground level O2 vmr of 10−6 PAL. Top: calculated with CABRRTM from paper I, bottom: calculated with CABMRACO2. Additionally, the total production rate of O2 (solid black line) is plotted.



Page 60 of 71

60

879 880 881 882 883 884 885 886 887 888 889 890 891 892

893 894 895

of O by diffusion from photolyzed CO2 in other atmospheric layers. Subclasses PA1, PB1 and PB2 increase by only a minor amount, i.e. by about 0.8, 0.8 and 0.4% (CABMRACO2), respectively, compared to paper I (CABRRTM) whereas PA2 decreases by 2.1%. The slight increase in PA1 results from an increase in P1 related to increased HOx concentrations in the upper atmosphere despite lower CO2 photolysis and a decrease in P2 due to lower HOx concentrations in the lower and middle atmosphere. The decrease in PA2 is exclusively related to the decrease in P4 as a result of a weaker CO2 photolysis rate which leads to a decrease in O. Fig. 34 illustrates the resulting changes in the altitude dependence of the O2 production pathways. In the present work (lower panel) the total column-integrated O2 production rate is decreased by 18% especially in the middle and lower atmosphere due to decreases in the O2 production pathways P2 and P4 in comparison to paper I (upper panel) because of less efficient production of O from CO2 photolysis. The O2 destruction pathways can be categorized into 4 classes: • class LA: O2 is destroyed by pathways with the net reaction O2 + 2CO → 2CO2 (COoxidation) which are catalyzed predominantly1 by HOx species

896

– subclass LA1 (L1, L2, L6, L8): without O2 photolysis,

897

– subclass LA2 (L3, L5): initiated by the photolysis of O2 ,

898 899

• class LB: O2 is consumed by CH4 -oxidation pathways

900

– subclass LB1 (L4, L7): net reaction 3O2 + 2CH4 → 4H2 O + 2CO,

901

– subclass LB2 (L10): net reaction O2 + CH4 → H2 O + H2 + CO,

902 903 904

905 906 907

908 909 910 911

• class LC (L9): O2 is consumed by pathways with the net reaction O2 + 2H2 → 2H2 O (H2 oxidation), • class LD (L11): O2 is consumed by pathways with the net reaction O2 + H2 O + CO → H2 O2 + CO2 . For O2 destruction classes and pathways only minor changes are observed. Fig. 35 illustrates the resulting changes in the altitude dependence of the O2 destruction pathways. In the present work (lower panel) for the total columnintegrated O2 destruction rate a decrease by 19% is observed compared to paper 1

L6 is also catalyzed by NOx .


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61

Fig. 35: Same notation as for Fig. 34 but for destruction of O2 . Note, the destruction of O2 is composed of a large number of pathways individually contributing less than 1.5% (white area).

914

I due to less CO in the lower atmosphere (see Fig. 33). This was due to lower photolytic production of CO from H2 CO in this region of the atmosphere due to less UV radiation.

915

A.2.3 Biogeochemical Module

912 913

916 917 918 919 920 921

Paper I identified two regimes with respect to the NPP - firstly a low O2 regime prior to the GOE and secondly a high O2 regime after the GOE. It was shown that in the low O2 regime, N is sensitive to the governing atmospheric chemistry, i.e. the net chemical change of O2 . In the case of Archean Earth, the resulting value N for the NPP decreases by 21% compared to paper I due to the strong change in the column-integrated net chemical change (P − L).



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936 937 938 939

940 941 942 943

944 945 946

947 948 949

950 951

952 953 954

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