Global tropospheric halogen (Cl, Br, I) chemistry and its impact on oxidants

. We present an updated mechanism for tropospheric halogen (Cl + Br + I) chemistry in the GEOS-Chem global atmospheric chemical transport model and apply it to investigate halogen radical cycling and implications for tropospheric oxidants. Improved representation of HOBr heterogeneous chemistry and its pH dependence in our simulation leads to less effective recycling and mobilization of bromine radicals, and enables the model to include mechanistic sea salt aerosol 25 debromination without generating excessive BrO. The resulting global mean tropospheric BrO mixing ratio is 0.19 ppt, lower than previous versions of GEOS-Chem. Model BrO shows variable consistency and biases in comparison to surface and aircraft observations in marine air, which are often near or below the detection limit. The model underestimates the daytime measurements of Cl 2 and BrCl from the ATom aircraft campaign over the Pacific and Atlantic, which if correct would imply a very large missing primary source of chlorine radicals. Model IO is highest in the marine boundary layer and uniform in the free troposphere, with a 30 global mean tropospheric mixing ratio of 0.08 ppt, and shows consistency with surface and aircraft observations. The modeled global mean tropospheric concentration of Cl atoms is 630 cm -3 , contributing 0.8% of the global oxidation of methane, 14% of ethane, 8% of propane, and 7% of higher alkanes. Halogen chemistry decreases the global tropospheric burden of ozone by 11%, NO x by 6%, and OH by 4%. Most of the ozone decrease is driven by iodine-catalyzed loss. The resulting GEOS-Chem ozone simulation is unbiased in the Southern Hemisphere but too low in the Northern Hemisphere.

Reaction of Clˉ with N2O5 in polluted environments at night produces ClNO2 that photolyzes in the daytime to return Cl atoms and NO2, stimulating ozone production Roberts et al., 2008). Acid displacement of Clˉ by HNO3 is a source of NO3ˉ aerosol. Reviews by Saiz- Lopez and von Glasow (2012) and Simpson et al. (2015) describe this fundamental knowledge of 45 tropospheric halogen chemistry in more detail.
A number of global modelling studies have explored the importance of halogen chemistry in the troposphere (von Glasow et al., 2004;Saiz-Lopez et al., 2006;Ordóñez et al., 2012;Long et al., 2014), but there remain large uncertainties in sources and chemical mechanisms. Here we present a new mechanistic description of halogen tropospheric chemistry in the GEOS-Chem global model 50 that synthesizes previous GEOS-Chem developments (Parrella et al., 2012;Eastham et al., 2014;Schmidt et al., 2016;Sherwen et al., 2016a;Sherwen et al., 2017;Chen et al., 2017;Wang et al., 2019;Zhu et al., 2019) and includes a number of updates. We use the updated model to interpret recent observations of tropospheric halogens, describe halogen radical cycling, and quantify the impacts on tropospheric oxidant chemistry. Shah et al. (2021) examines the impact of our simulated Br and Cl atom concentrations in a new redox mechanism for atmospheric mercury. 55

Tropospheric halogen chemistry in GEOS-Chem
We describe here our updated representation of tropospheric halogen chemistry in version 12.9 of GEOS-Chem (http://www.geoschem.org), implemented as part of the general model mechanism for coupled ozone-NOx-VOCs-aerosol-halogen tropospheric and stratospheric chemistry. Extensive referencing will be made to Sherwen et al. (2016b), who implemented the previous representation of tropospheric halogen chemistry in , and to Wang et al. (2019), who described an 60 earlier version of the mechanism implemented here. GEOS-Chem stratospheric halogen chemistry is as described by Eastham et al. (2014) and we will not discuss it further here. Table 1 lists the global sources and sinks of tropospheric gas-phase inorganic chlorine (Cly), bromine (Bry), and Iodine (Iy) in GEOS-Chem. SSA emissions are from Jaeglé et al. (2011). Open fire emissions of HCl are obtained by applying the emission 65 factors from Andreae (2019) for different vegetation types to the GFED4 (Global Fire Emissions Database version 4) biomass burned inventory (van der Werf et al., 2017). The resulting global source of 0.5 Tg Cl a -1 is much smaller than in Wang et al. (2019), who used older emission factors from Lobert et al. (1999). Fuel combustion, waste incineration, and dust are additional sources of HCl and Clˉ but the emissions are highly uncertain and likely negligible from a global budget perspective (Wang et al., 2019). We do not include them here. Organohalogen gases can produce halogen radicals by oxidation and photolysis. Emissions 70 of long-lived organohalogen gases (CH2Cl, CH2Cl2, CHCl3, CHBr3) are implicitly treated in the model by specifying latitudinally and monthly surface air boundary conditions from CMIP6 (Historical greenhouse gas concentrations for climate modelling) (Meinshausen et al., 2017). Emissions of short-lived bromocarbons (CH3Br, CH2Br2) and iodocarbons (CH3I, CH2I2, CH2ICl, CH2IBr) are from Bell et al. (2002), Liang et al. (2010), and Ordóñez et al. (2012). The main global source of tropospheric Cly is mobilization of Clˉ from SSA. A total of 50 Tg Clˉ a -1 (2.4% of SSA emissions) is mobilized to Cly in the model by acid displacement and other heterogeneous reactions. This number is smaller than our previous estimate in Wang et al. (2019) (64 Tg Clˉ a -1 ), mainly due to slower ClNO2 generation from the N2O5 + Clˉ reaction (Section 2.3 and Text S1)). Organochlorines provide a tropospheric source of 3.3 Tg Clˉ a -1 as Cl atoms from photolysis and oxidation. Transport from stratosphere adds 0.14 Tg Cl a -1 to tropospheric Cly. The source of Iy is estimated to be 2.7 Tg I a -1 , mostly from the inorganic 80 iodine (HOI, I2) formed from the ocean surface reaction of O3 with iodide (I − ), based on Carpenter et al. (2013) and MacDonald et al. (2014), and as described by Sherwen et al. (2016b).

Sources of tropospheric halogens
In GEOS-Chem versions before 12.9, SSA debromination was not included despite being known to be an important source for Bry (Sander et al., 2003). This is because SSA debromination generated excessive BrO concentrations in comparison to observations, 85 which then drove excessive ozone depletion (Schmidt et al., 2016;Zhu et al., 2019). Revision of HOBr reactive uptake as source of bromine radicals effectively corrects this problem (Section 2.2), allowing us to include mechanistically the SSA debromination source. This provides the main global source of tropospheric Bry (20 Tg Br a -1 ), mostly through the HOBr/HOCl/HOI + Brˉ heterogeneous reactions. Bromocarbon gases contribute only 0.54 Gg Br a -1 to Bry but still dominate the Bry source in the free troposphere. 90

Chemical mechanism
Our tropospheric halogen chemistry mechanism synthesizes and updates previous GEOS-Chem mechanistic developments.
Tropospheric bromine chemistry was first built by Parrella et al. (2012) with updates to heterogeneous reactions by Schmidt et al. 95 (2016), Chen et al. (2017), Wang et al. (2019), andZhu et al. (2019). Iodine chemistry was built by Sherwen et al. (2016a) and Sherwen et al. (2016b). Recent general model updates important for halogen chemistry include a new method of simulating cloud chemistry in partly cloudy grid cells that accounts for limitation by entrainment of air into the cloud (Holmes et al., 2019) and an improved cloudwater pH calculation that considers carboxylic acids and dust alkalinity (Moch et al., 2020;Shah et al., 2020).
We update here the reactive uptake of HOBr by aerosols and cloud droplets ( where Y is the yield of Br2 and 1-Y is the yield of BrCl, which are calculated based on the laboratory study of Fickert et al. (1999) and described in Table 2.
Total reactive uptake of HOBr from reactions (R3)-(R5) in aqueous aerosols and clouds is calculated with a standard first-order reactive uptake coefficient γ (Jacob, 2000), calculated following Ammann et al. (2013): 115 where HHOBr is the the Henry's law constant of HOBr, T is temperature, R is the universal gas constant; Dl is the liquid phase 120 diffusion coefficient for HOBr; f(r, Ir) is the reacto-diffusive correction term, and k I is the first-order total reaction rate constant of HOBr from pathways (R3-R5) computed as a function of the concentrations of Brˉ, Clˉ, H + , HSO3ˉ, and SO3 2ˉ. After computing the overall loss of HOBr, we distribute the loss by pathways on the basis of the relative reaction rates . Reactions (R3) and (R4) are important only in clouds because dissolution of SO2 depends on the liquid water content. Wang et al. (2019) previously calculated 5 based on experimental results over limited and inconsistent pH ranges (pH = 1.9-2.4 125 for HOBr+Brˉ, pH = 6.4 for HOBr+Clˉ (Beckwith et al., 1996;Liu and Margerum, 2001)). This generated excessive BrO concentrations in comparison to observations. Here we revise the calculation of 5 to consider the entire range of aerosol and cloud pH, as recommended by Roberts et al. (2014), resulting in much slower rate. We also adopt a new value for 3 from a recent laboratory study (Liu and Abbatt, 2020). Details of these updates are in Table 2. The overall result is to have less efficient heterogeneous recycling and mobilization of bromine radicals in both aerosols and clouds. Aerosol aqueous-phase reaction of N2O5 with Clˉ produces ClNO2 that photolyzes in the daytime to return Cl atoms and NO2. The 140 reaction competes with N2O5 hydrolysis, with the following first-order loss representation for N2O5: and recommended lower values than previously used in GEOS-Chem by Wang et al. (2019) to account for the effect of organic coating of particles. We previously implemented this update in Wang et al. (2020) and it is now part of GEOS-Chem version 12.9. 145 Additional updates to the GEOS-Chem halogen mechanism in version 12.9 include a new scheme to calculate the reactive uptake coefficients γ on ice crystals following recommendations by the International Union of Pure and Applied Chemistry (IUPAC) (Crowley et al., 2010) as listed in Table 3. We calculate the effective radius of ice crystals based on air temperature following Heymsfield et al. (2014) and Holmes et al. (2019), and increase the resulting surface area by a factor of 2.25 to account for irregular 150 shape (Schmitt and Heymsfield, 2005). We also include the temperature dependence for BrNO3 hydrolysis in Deiber et al. (2004) (Text S1), and minor updates for HOI, IONO, and IONO2 heterogeneous reactions (Text S2). Figure 1 shows the global budgets and cycling of tropospheric inorganic chlorine (1a), bromine (1b), and iodine (1c) in our model simulation. Figure 2 shows the annual mean global distributions of Cl atoms, BrO, and IO. Figure 3 shows the global mean vertical 155 distribution of the halogen speciation for reactive chlorine (Cl* ≡ Cly -HCl), Bry, and Iy. GEOS-Chem is driven here by 2016 GEOS-FP (forward processing) assimilated meteorological fields from the NASA Global Modeling and Assimilation Office (GMAO) with native horizontal resolution of 0.25° x 0.3125° and 72 vertical levels from the surface to the mesosphere. Our model simulation is conducted at 4° x 5° horizontal resolution and meteorological fields are conservatively degraded to that resolution.

Global budget and distribution of tropospheric halogens
The model is spun up for 1 year for initialization. 160

Chlorine
The dominant global source of Cly is acid displacement from SSA to HCl. The global rate of HCl generation from acid displacement is 46 Tg Cl a -1 and close to the observationally based estimate of 50 Tg Cl a -1 by Graedel and Keene (1995). HCl is the largest reservoir of tropospheric Cly, with a global mean tropospheric mixing ratio of 45 ppt. Most of HCl is removed by deposition, and only a small fraction (7.3 Tg Cl a -1 ) reacts with OH and contributes to reactive chlorine Cl* (≡ Cly -HCl). Cl* can be also generated 165 from Clˉ in clouds and aerosols by heterogeneous reactions with principal contributions from HOBr+Clˉ (2.6 Tg Cl a -1 ), HOCl+Clˉ (1.5 Tg Cl a -1 ), HOI/IONOx + Clˉ (0.8 Tg Cl a -1 ), and N2O5+Clˉ (0.68 Tg Cl a -1 ). This heterogeneous source of 6.3 Tg Cl a -1 is lower than our previous estimate of 12 Tg Cl a -1 in Wang et al. (2019), since the updated mechanisms for HOBr+Clˉ (Section 2.3) and N2O5+Clˉ (Section 2.4 and Wang et al. (2020)) reactions are slower. We calculate a tropospheric lifetime of 2.3 hours for Cl*.
Loss of Cl* is mainly through the reaction of Cl with methane (44%) and other organic compounds. 170 Distributions of Cl* in the troposphere are generally similar to Wang et al. (2019). As shown in Figure 2, tropospheric Cl atom concentrations are highest at the surface, reflecting the source from SSA, and in the upper troposphere due to transport from the stratosphere as well as cold temperature slowing down the Cl + methane reaction. In surface air, Cl atom concentrations are usually highest along polluted coastlines where the large sources of HNO3, H2SO4, and N2O5 from anthropogenic emissions drive acid 175 displacement and ClNO2 production. Figure 3 shows the global zonal mean vertical distribution of Cl* species. Boundary layer Cl* is dominated on a zonal mean basis by ClNO2 formed from N2O5+Clˉ in polluted air. High mixing ratios of ClNO3 in the upper troposphere are related to transport from the stratosphere and its slow hydrolysis. The BrCl mixing ratio is much lower than in the

Bromine
The largest source of Bry is from SSA debromination in the marine boundary layer (MBL), mainly contributed by HOBr+Brˉ and O3+Brˉ producing Br2 and HOBr, respectively. Bromocarbon photochemistry dominates the source of Bry in the free troposphere.
Uptake of HBr by SSA is the major sink of Bry. The global tropospheric loading of BrO in the model is 2.1 Gg Br, corresponding to a mean tropospheric mixing ratio of 0.19 ppt (0.38 ppt in daytime). This value is much lower than the most recent GEOS-Chem 185 estimate of 8.0 Gg by Zhu et al. (2019), because of the updated HOBr heterogeneous chemistry described in Section 2.3. The newly added pH-dependences in Table 2 decrease the rate of reaction (R5), resulting in much slower recycling of HOBr in cloud and aerosol water. HOBr is now more likely to react with S(IV) via reactions (R3) and (R4), forming HBr which then gets taken up by SSA. In Zhu et al. (2019), 82% of HOBr heterogeneous reactions were with Brˉ and Clˉ, and only 18% were with S(IV). Due to the update in Section 2.3, 59% of HOBr heterogeneous reactions are with Brˉ and Clˉ, and 41% are with S(IV). The higher fraction 190 of Bry as HBr decreases the tropospheric lifetime of Bry because HBr is more water-soluble than other Bry species. We calculate tropospheric lifetimes of 7.9 hours for Bry and 6.8 minutes for BrOx (≡ Br+BrO). Figure (Table 3). Figure 3 shows the global mean vertical distribution of Bry species, which is very different from Sherwen et al. (2016b) where the Bry concentration increased with altitude. This is due to the inclusion of SSA debromination in our simulation. Our Bry mixing ratio in the MBL is still only slightly higher than that in Sherwen et al. (2016b) because of the 200 much lower Bry lifetime resulting from the slower HOBr heterogeneous reactions, as mentioned above.

Iodine
The Iy source totals 2.7 Tg I a -1 with most (2.1 Tg I a -1 ) originating from ocean volatilization of HOI and I2 (Carpenter et al., 2013;MacDonald et al., 2014). The sink of Iy is from deposition (1.8 Tg I a -1 ) and uptake by aerosols (0.91 Tg I a -1 ). The global tropospheric loading of IO in the model is 1.4 Gg I, corresponding to a mean tropospheric mixing ratio of 0.08 ppt. As shown in 205 Surface IO mixing ratios are highest over tropical oceans, where both organic and inorganic iodine emissions are high due to the high temperature. Concentrations of IO and most Iy species are the lowest in middle troposphere where Iy speciation is mostly as HOI, which can be removed via wet deposition efficiently. IO is higher in the upper troposphere where its cycling is mainly with IONO2 and wet deposition is less efficient. We calculate tropospheric lifetimes of 1.6 days for Iy and 1.7 minutes for I+IO* (≡ 210 IO+OIO+2I2O2+2I2O3+2I2O4). Our results are consistent with Sherwen et al. (2016b) since the iodine chemistry is largely unchanged. Our only significant update has been to conserve mass in iodine heterogeneous reactions (Text S2) but this has little impact.

Comparison to observations
Here we compare the model simulation for 2016 to observations for gas-phase halogen species collected from surface and aircraft 215 campaigns. The observations are in different years but we assume that interannual variability is small compared to other sources  (Table 4). There is no evidence of systematic model bias but more sensitive observations would be needed to be conclusive. The WINTER aircraft campaign provided data for multiple Cly gases including HCl, ClNO2, HOCl, and Cl2. The measurements were made over the eastern US and offshore during February-March 2015 by I-TOF-CIMS , as summarized in 255 Table 4. Figure 8 compares the observed median vertical profiles of HCl, ClNO2, HOCl, and Cl2 during WINTER to the model sampled along the flight tracks for the corresponding period. Modeled HCl is lower than the observations but mostly within the calibration uncertainty (± 30%). Modeled HOCl largely underestimates WINTER observations. Wang et al. (2019) found that such underestimation is over both land and ocean and mainly in daytime when HOCl has very short lifetime against photolysis (a few minutes). This may suggest a large photochemical source needed to decrease the model bias. Recent work also identified to the 260 potential for IOx-ion chemistry to lead to measurement interferences (Dörich et al., 2021), of the detection of acid gases which could impact the measured HOCl:HCl ratio. Furthermore, rapid interconversion of halogen species on inlet walls have been reported that could also impact the measured HOCl:HCl ratio (Neuman et al., 2010).  (Table 4). The underestimates of HOCl during WINTER, and BrCl, Cl2 during ATom at daytime may suggest a large photochemical source that can produce chlorine radicals from Clˉ.

Global implications for tropospheric oxidant chemistry
We now examine the implications of tropospheric halogen chemistry as described by our mechanism on the concentrations of tropospheric VOCs, ozone, NOx, and OH. Shah et al. (2021) examined the implications for mercury chemistry.

Volatile organic compounds (VOCs)
Cl atoms are strong VOC oxidants, but their importance is limited by their small supply. The global mean tropospheric Cl atom 285 concentration in our model is 630 cm -3 , consistent with the upper limit of 1000 cm -3 inferred by Singh et al. (1996) Figure 12 shows the effects of halogen chemistry on tropospheric OH, NOx, and ozone concentrations, as obtained by difference 295 with a sensitivity simulation excluding all halogen reactions in the troposphere ("no halogen"). Halogen chemistry decreases the global tropospheric ozone burden by 11% in our model, which is smaller than the 18.6% in Sherwen et al. (2016b). Global ozone chemical production decreases by 2% while ozone lifetime decreases by 10%. The decrease in ozone production is due to a 5.6% global decrease in NOx as a result of formation and hydrolysis of halogen nitrates XNO3 (X ≡ Cl, Br, I):

NO 3 + H 2 O → HO + HNO 3 (R18)
Globally, such NOx loss is mostly through ClNO3 and BrNO3 hydrolysis, with negligible contribution from INO3. As shown in Figure 12, surface NOx increases over the continents and this is due to ClNO2 chemistry. We previously showed in Wang et al.
(2019) that Clˉ originating from SSA can be transported far inland by acid displacement of HCl and subsequent HCl uptake by sulfate-nitrate-ammonium (SNA) aerosols. Halogen chemistry in our model lowers global tropospheric concentrations of OH and 305 HO2 by 4.1% and 3.4% respectively. Decrease in OH is mainly due to the decrease of ozone, which reduces primary OH production from ozone by 9.8%. The increase in OH over continental regions (Figure 12) is due to ClNO2 chemistry. Table 5 summarizes the global annual budget of tropospheric ozone in the standard model and in the no halogen simulation. The budget of ozone is shown as that of odd oxygen (Ox ≡ O3 + O + O( 1 D) + NO2 + 2NO3 + peroxyacylnitrates + HNO3 + HNO4 + 310 3N2O5 + organic nitrates + Criegee intermediates + XO + HOX + XNO2 + 2XNO3 + 2OIO + 2I2O2 + 3I2O3 + 4I2O4 + 2Cl2O2 + 2OClO where X ≡ Cl, Br, I) to account for the rapid cycling between Ox species. The 11% shorter ozone lifetime as a result of halogen chemistry is due to catalytic ozone loss cycles driven by iodine (7.6%), bromine (2.6%) and chlorine (0.3%). Figure 13 shows the relative contributions of different reaction routes to ozone chemical loss in troposphere. Halogens contribute about 19% https://doi.org/10.5194/acp-2021-441 Preprint. Discussion started: 2 June 2021 c Author(s) 2021. CC BY 4.0 License. of ozone loss in the MBL, decreasing to 8% at 2-4 km altitude and then increasing to 24% in the upper troposphere. Halogen-315 catalyzed ozone loss is mainly driven by the sequence (X ≡ I, Br, Cl): Bates and Jacob (2020) introduced an expanded odd oxygen family, Oy≡Ox+Oz, to include both Ox and an additional subfamily, 320 Oz, consisting of HOx and its reservoirs (Oz ≡ 0.5×(H + OH + organic peroxy radicals + HNO2 + HNO3 + HNO4 + peroxyacylnitrates + organic nitrates + X + XO + XNO2 + XNO3 + OIO + OClO + ClOO) + H2O2 + organic peroxides + X2 + HOX + I2O2 +I2O3 + I2O4 + Cl2O2 where X ≡ Cl, Br, I). Table 4 also summarizes the budget of Oz. The global tropospheric Oz burden decreases by 4% due to the halogen chemistry, which is mainly because of the lower production from Ox. Following Bates and Jacob (2020), we define the chain length N, or Ox production efficiency per unit Oz, as the number of times a unit of Oz is 325 converted to Ox before it is removed to terminal sinks: In the conventional Ox budget analysis, conversion from Ox to Oz through O( 1 D) + H2O, is viewed as a sink for Ox; but if N > 1 it is actually a net source. By considering this, Bates and Jacob (2020) introduced an effective ozone lifetime as: where ki is the pseudofirst-order loss rate constant for process i. As shown in Table 4, N increases from 1.40 to 1.47 by including halogen chemistry, thus amplifying ozone production efficiency from O( 1 D) + H2O. This is because of the decrease of HO2 which slows down the loss rate of HOx. The effective ozone lifetime decreases by 15%, from 71 to 60 days, because the halogen-driven catalytic pathways represent true ozone sinks by converting O3 to O2.  Table S1) and we average the data into six latitudinal bands. Halogen chemistry does not degrade the simulation in the Southern Hemisphere, where the model bias is small, but worsens the underestimate in the Northern Hemisphere. Similar 340 results are found in Figure 14, which compares modeled surface ozone mixing ratios to observations at surface sites. There is no significant seasonal variation for the impacts of halogen chemistry on surface ozone at these sites. The last extensive evaluation of the global tropospheric ozone simulation in GEOS-Chem was done by Hu et al. (2017) and found no significant bias, but it used

Conclusions
We presented a new comprehensive representation of tropospheric halogen chemistry in the GEOS-Chem model that synthesizes and updates previous model developments. We used it to analyze the sources and cycling of halogen radicals, evaluate against observations of halogen radicals and their reservoirs, and examine the implications for tropospheric oxidant chemistry. 350 The model includes an improved representation of heterogeneous chemistry in aerosols and clouds, including in particular the reactions of HOBr, leading to less effective recycling and mobilization of bromine radicals. This allows us to include in the model the known source of bromine radicals from debromination of sea salt aerosols (SSA) without generating excessive BrO concentrations. Simulation of cloud processing is improved to include a more accurate computation of cloudwater pH (Shah et al., 355 2020) and cloud entrainment (Holmes et al., 2019). ClNO2 production by the heterogeneous N2O5 + Clˉ reaction is updated to a slower rate to account for organic coating of particles (McDuffie et al., 2018a;. Cycling of chlorine and iodine radicals is similar to previous versions of GEOS-Chem (Wang et al., 2019;Sherwen et al., 2016) but cycling of bromine radicals is very different. We find a mean tropospheric BrO mixing ratio of 0.19 ppt, much lower than 360 previous GEOS-Chem estimates and reflecting the less effective heterogeneous recycling of bromine radicals. BrO is highest in the marine boundary layer (MBL) where SSA debromination is the main source, and in the upper troposphere, due to photodecomposition of bromocarbons and transport from the stratosphere. Model results are consistent with MBL observations of BrO from coastal sites and ship cruises, though observations are often below the detection limit. Comparisons to vertical profiles from aircraft campaigns paints an inconsistent picture, with model BrO being lower than the CAST CIMS, CONTRAST DOAS, 365 and TORERO DOAS measurements over the tropical Pacific, but higher than the ATom CIMS measurements at high altitudes on Pacific and Atlantic transects. The TORERO and CONTRAST DOAS data show increasing BrO concentrations in the upper troposphere but the ATom CIMS data do not. The aircraft observations are again below or close to detection limits. A more confident evaluation of tropospheric bromine radical chemistry will require more sensitive observations of BrO and its reservoirs in the future. 370 Our simulation shows a global mass-weighted mean Cl atom concentration of 630 molecules cm -3 in the troposphere. Oxidation by Cl atoms accounts for 0.8% of the global loss of atmospheric methane and has larger effects on the global losses of ethane (14%), propane (8%), and higher alkanes (7%). Reactive chlorine (Cl* ≡ Cly -HCl) is mainly generated from HCl + OH (7.3 Tg Cl a -1 ), heterogeneous reactions of Clˉ in clouds (6 Tg Cl a -1 ) and oxidation of organochlorines (3.3 Tg Cl a -1 ). Comparisons of 375 model results to observations in marine surface air and aircraft campaigns in this study and our previous work (Wang et al., 2019) show that the model is in general consistent with the range and distributions of observed HCl and ClNO2 concentrations. Halogen chemistry decreases the global burden of tropospheric ozone in GEOS-Chem by 11%. This reflects a 2% decrease in ozone production due to the sink of NOx from formation and hydrolysis of ClNO3 and BrNO3, and a 11% increase in ozone 385 chemical loss due to catalytic cycles involving iodine (8%) and bromine (3%). The global mean tropospheric OH concentration decreases by 4.1%, mostly due to the decrease in ozone. Tropospheric ozone concentrations in GEOS-Chem show no significant bias in the Southern Hemisphere relative to ozonesonde data, but a low bias in the Northern Hemisphere that is also present in the absence of halogen chemistry. Addressing this low bias should be a priority for future research.
Data availability. The model code is available at GEOS-Chem repository (http://www.geos-chem.org). ATom data are publicly 390 available through the Oak Ridge National Laboratory DAAC (https://daac.ornl.gov/ATOM/campaign/data.html). Data from the WINTER, CONTRAST, and TORERO campaigns are publicly available at the EOL data archive (https://data.eol.ucar.edu/). Data of CAST are publicly available at the CEDA archive (https://catalogue.ceda.ac.uk/uuid/565b6bb5a0535b438ad2fae4c852e1b3).
All links mentioned here were last accessed on 1 April 2021.  Table 1. Global sources and sinks of tropospheric gas-phase phase inorganic chlorine (Cly), bromine (Bry), and iodine (Iy) a .