Measurement and modelling of reactive halogen species over the tropical Atlantic Ocean

Measurement and modelling of reactive halogen species over the tropical Atlantic Ocean A. S. Mahajan, J. M. C. Plane, H. Oetjen, L. Mendes, R. W. Saunders, A. Saiz-Lopez, C. E. Jones, L. J. Carpenter, and G. B. McFiggans School of Chemistry, University of Leeds, Leeds LS2 9JT, UK Instituto Nacional de Meteorologia Geofisica (INMG), Delegação de São Vicente, Monte, CP 15, Mindelo, Cape Verde Laboratory of Atmospheric and Climate Science, Consejo Superior de Investigaciones Cientificas, 45007 Toledo, Spain Department of Chemistry, University of York, York, UK Centre for Atmospheric Science, School of Earth, Environment and Environmental Sciences, The University of Manchester, Manchester M13 9PL, UK

iodine monoxide (IO) and bromine oxide (BrO) over eight months in the tropical open ocean MBL, on the north-eastern side of S ão Vicente (Cape Verde Islands, 16.85 • N, 24.87 • W).The highest BrO mixing ratio observed was 5.6±1 ppt, while the maximum observed IO mixing ratio was 3.1±0.4ppt.The average values seen between 09:00-17:00 GMT were ∼2.8 ppt for BrO and ∼1.5 ppt for IO; these averages showed little variability over the entire campaign from November 2006 to June 2007.A 1-dimensional chemistry and transport model is used to study the evolution of iodine species and quantify the combined impact of iodine and bromine chemistry on the oxidising capacity of the MBL.It appears that the measured fluxes of iodocarbons are insufficient to account for the observed levels of IO, and that an additional I atom source is required, possibly caused by the deposition of O 3 onto the ocean surface in the presence of solar radiation.Modelling results also show that the total O 3 depletion observed at Cape Verde cannot be explained in the absence of halogen chemistry, which contributes ∼45% of the total O 3 depletion at the height of measurements (10 m) during summer.
The model also predicts that halogens decrease the hydroperoxy radical (HO 2 ) concentration by ∼14% and increase the hydroxyl radical (OH) concentration by ∼13% near the ocean surface.The oxidation of dimethyl sulphide (DMS) by BrO takes place at a comparable rate to oxidation by OH in this environment.Finally, the potential of iodine chemistry to form new particles is explored and conditions under which particle formation could be important in the remote MBL are discussed.Introduction

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Relatively high BrO mixing ratios (>10 ppt) are consistently reported in the springtime polar BL (e.g., Hausmann and Platt, 1994;Tuckermann et al., 1997;Richter et al., 1998;Wagner and Platt, 1998;Saiz-Lopez et al., 2007a).BrO has also been measured on a regional scale above salt lakes and in volcanic plumes (Hebestreit et al., 1999;Stutz et al., 2002;Bobrowski et al., 2003).In the mid-latitude MBL, BrO has been positively detected only at three coastal locations: Leser et al. (2003) observed variable levels around the Canary Islands peaking at 2.4 ppt, Saiz-Lopez et al. (2004) saw post sunrise peak levels of ∼6 ppt with an average of 2 ppt over the day at Mace Head, Ireland, while Mahajan et al. (2009a) reported elevated BrO mixing ratios peaking at 7.5 ppt in the semi-polluted coastal environment at Roscoff, France.Modelling studies have predicted background BrO levels of 1-4 ppt (Vogt et al., 1996;von Glasow et al., 2002a).Upper limits have also been reported where BrO was not seen above the instrumental detection limit.Theys et al. (2007) reported BrO in the free troposphere (FT) around the Reunion Island with an upper limit of 0.5 ppt in the BL.H önninger (2002) also calculated that BrO around Crete was less than 1.5 ppt, while Peters et al. (2005) reported an upper limit of 2 ppt from Lilia, on the coast of Brittany, France.
IO was first detected at Mace Head, Ireland by Alicke et al. (1999).Since then, measurements at a variety of sub-tropical and temperate coastal locations have reported IO levels up to ∼27 ppt (e.g., Allan et al., 2000;Frieß et al., 2001;Saiz-Lopez and Plane, 2004;Zingler and Platt, 2005;Peters et al., 2005;Stutz et al., 2007;Whalley et al., 2007;Mahajan et al., 2009b).Upper limits for IO have been reported at other locations: 0.8 ppt at Crete (H önninger, 2002), 2 ppt at Ny-Ålysund (Wittrock et al., 2000), and 2 ppt at Kuujjuarapik, Canada (H önninger et al., 2004).In contrast, IO has been observed in excess of 20 ppt in coastal Antarctica during springtime (Saiz-Lopez et al 2007a), and satellite measurements have shown that IO can be detected over large areas in and around the continent (Saiz-Lopez et al., 2007b;Sch önhardt et al., 2008).Observations of other iodine compounds such as I 2 and OIO are not as common as BrO or IO.OIO was first detected at Cape Grim (Tasmania) by Allan et al. (2001), while the first measurement of I 2 was made by Saiz-Lopez and Plane (2004) at Mace Head.
A recent study at the Cape Verde Islands reported the presence of both BrO and IO at mixing ratios of 1-3 ppt, and indicated that the observed ozone depletion during the same campaign could not be explained in the absence of halogen compounds (Read et al., 2008).The importance of this study is that this was arguably the first measurement of these halogen radicals in the mid-ocean MBL where local biological sources (such as macroalgae exposed at low tide) were unimportant, and hence may be representative of halogen activity over much of the open ocean (Read et al., 2008).
In the present paper, we consider these halogen oxide measurements at Cape Verde in further detail, using a 1-D model to investigate the source strengths and likely vertical distributions of iodine species in the tropical Atlantic MBL.We then explore the effects of combined iodine and bromine chemistry on O 3 depletion, the HO x ratio and the lifetime of DMS.Finally, we consider the potential for iodine oxides to form new particles which may grow to become cloud condensation nuclei (CCN) in this environment.the north-eastern side of the island of S ão Vicente.Further details about the observatory and the prevalent meteorological conditions can be found elsewhere (Lee et al., 2009).The long-path DOAS instrument (LP-DOAS) was placed in a temperaturecontrolled sea freight container at the observatory.The optical path ran across the bay, where a retro-reflector array was placed at a distance of 6 km from the measurement site (Supporting online text Fig. S1, http://www.atmos-chem-phys-discuss.net/ 9/24281/2009/acpd-9-24281-2009-supplement.pdf),resulting in a total optical path length of 12 km.The light beam was ∼10 m above the ocean surface.Further details on the LP-DOAS and the deconvolution procedure used can be found elsewhere (Plane and Saiz-Lopez, 2006).Briefly, the recorded spectra were converted into optical densities, and the contributions of the individual absorbing species were then determined by using singular value decomposition to fit simultaneously a library of reference absorption cross-sections.For the retrieval of BrO, the 324-357 nm spectral window was used, while IO was retrieved in the 425-445 nm window.Reference spectra included in the fitting routine for retrieving BrO were: BrO (Wilmouth et al., 1999), NO 2 (Vandaele et al., 1998), O 3 (Daumont et al., 1992), HONO (Bongartz et al., 1991) and HCHO (Cantrell et al., 1990).For retrieving IO, reference spectra were: IO (G ómez Martin et al., 2005), NO 2 (Harder et al., 1997) and H 2 O (Rothman et al., 2003).An example of the quality of the fits, along with averaged diurnal and annual measurements are presented elsewhere (Read et al., 2008).
During the summer of 2007 (May), an intensive campaign took place at the observatory as a part of the Reactive Halogens in the Marine Boundary Layer (RHaMBLe) project.The details of the instruments deployed during this period can be found in Lee et al. (2009), this issue.Along with other ancillary measurements, the intensive campaign also involved a research aircraft that measured composition over the site to assess how representative of the overlying BL it was and to determine any diurnal variability in BL depth.Additionally, a ship-based study measured the fluxes of halocarbons above the ocean surface around Cape Verde and over the upwelling region near Mauritania during the same period (Jones et al., 2009) spectively, for the eight months of the campaign.BrO and IO were observed above the detection limit whenever the solar zenith angle was less than ∼92 • , i.e., twilight and full daylight conditions.Although there was considerable day-to-day variation throughout the campaign, it can be seen that the profiles exhibit a "top-hat" distribution over the daytime with the maximum values between 09:00-17:00 GMT.On averaging the data over the entire campaign, both species show similar diurnal profiles dependent on solar radiation, with mixing ratios below the detection limits (BrO: 0.5-0.8ppt, IO: 0.3-0.5 ppt) at night and average peak values of ∼2.8 ppt BrO and ∼1.5 ppt IO during the daytime (Read et al., 2008).The monthly daytime-averaged values between 09:00 and 17:00 GMT, and the maximum observed mixing ratios in each month of the campaign, are shown for both species in Fig. 1c.There is clearly very little seasonal variation: although there is some hint of slightly higher values during spring (February-May), this is not significant within the variability of the observations.The BrO and IO mixing ratios do not show a strong correlation with the origin of the air masses given by the air mass back trajectories, the wind direction or the wind speed.Figure 2 shows the variation of BrO and IO with wind speed.There is no significant correlation for either species: for [BrO] vs. wind speed, linear regression through the data has a slope of 0.21±0.04 and R 2 =0.02, while for [IO] vs. wind speed the slope was 0.12±0.06and R 2 =0.003.In the case of BrO, all measurements were made when the wind happened to be from the prevailing north-easterly direction, and hence a correlation analysis cannot be made.In the case of IO, although the wind direction did vary, there does not appear to be clear dependence on wind direction.
Finally, the LP-DOAS was also periodically configured to measure other iodine Introduction

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Full 4 Discussion and modelling

Bromine chemistry
The release of inorganic bromine from sea-salt aerosols and its impact on the remote MBL has been studied in some detail in previous work (Vogt et al., 1996;von Glasow et al., 2000avon Glasow et al., , b, 2004;;Sander et al., 2003).Some model studies for the acid-catalysed activation of bromine from sea-salt aerosol have predicted BrO concentrations in the MBL ranging from 1-4 ppt, with a diurnal top-hat variation (Vogt et al., 1996;Yang et al., 2005).This is due to the removal of BrO via the reaction with HO 2 to form HOBr, which peaks at midday and limits the build up of BrO.Box model calculations for conditions seen at Cape Verde show that to sustain an average [BrO] of ∼2.8 ppt, there needs to be a total gas-phase bromine loading of ∼10 ppt.Considering that the [Br − ]/[Cl − ] ratio in fresh aerosols is ∼1.5×10 −3 molmol −1 , reflecting the sea water ratio of 1:660 (Wilson, 1975), and assuming an increased salinity of 320 gl −1 due to evaporation of water from sea-salt aerosols, the maximum amount of bromide available in 20 µm 3 cm −3 of wet aerosol volume (Allan et al., 2009) would be ∼4 ppt.Hence, continuous replenishment of aerosols and depletion of bromide from these aerosols would explain the levels of BrO seen at Cape Verde.Br − depletion in aerosols can occur through acid-catalysed reactions, which are pH-dependent (Mozurkewich, 1995;Abbatt and Waschewsky, 1998;Fickert et al., 1999).Br − depletion in aerosols has been reported previously (Gabriel et al., 2002;Newberg et al., 2005) 2009), indicating that the observed BrO levels can satisfactorily be explained through emissions from sea-salt aerosol.Since the focus of the modelling study described below is iodine chemistry, we will therefore not explicitly model bromine aerosol chemistry here.Instead, the model will be constrained with the BrO observations, and then used to explore the iodine source required to replicate the IO measurements and to study the combined impact of iodine and bromine chemistry on O 3 , HO x and DMS chemistry.
[BrO] is set to peak at 2.8 ppt with a top hat profile according to observations at Cape Verde and is considered to be well mixed in the boundary layer due to the long lifetime of inorganic bromine in the gas phase, which is ∼10 d in the tropics (von Glasow et al., 2004).

Iodine chemistry
Measurements and modelling studies of IO over the open ocean are rare, and the impacts of iodine chemistry on the remote oceanic environment have not been studied in detail.Although the levels of IO reported here are in accord with previous measurements along the coasts of Tenerife and Tasmania (Allan et al., 2000), no studies have been performed in the tropical remote MBL and to our knowledge these are the first measurements of reactive iodine species (RIS) over the tropical open ocean.The presence of IO above the detection limit of the instrument (0.3-0.5 ppt) at Cape Verde is surprising as the sources of iodine are assumed to be mainly biogenic.Although the oceans around Cape Verde are biologically active (Fallowski, 1997), there are no large macroalgal beds around S ão Vicente due to its steep shelf.Hence, it is probable that these waters represent a large area of increased iodine release.A recent study by Martino et al. (2009)  The Tropospheric HAlogen chemistry MOdel (THAMO) is now employed to test whether the levels of IO seen at Cape Verde can be explained by the measured iodocarbon fluxes.Further details about THAMO can be found in Saiz-Lopez et al. (2008).Briefly, it is a 1-D chemistry transport model with 200 stacked boxes at a vertical resolution of 5 m (total height 1 km).The model treats iodine, bromine, O 3 , NO x and HO x chemistry using 210 reactions.The chemical scheme used for this study is from Saiz-Lopez et al. (2008) and is updated according to Mahajan et al. (2009b) (Read et al., 2008;Lee et al., 2009, this issue;Read et al., 2009).Wind speed measurements collected at three heights (4, 10 and 30 m) were used to construct an eddy diffusion coefficient (K z ) profile (Stull, 1988), shown in the supporting online text (Fig. S2, http://www.atmos-chem-phys-discuss.net/9/24281/2009/ acpd-9-24281-2009-supplement.pdf).The average background aerosol surface area (ASA) used is 1×10 −6 cm 2 cm −3 , an average value estimated from Allan et al. (2009).During the RHaMBLe campaign, aircraft measurements of temperature in the BL revealed a strong temperature inversion about 1 km from the surface (Read et al., 2008), indicating that the BL at Cape Verde is decoupled from the free troposphere.
Hence, the K z profile is assumed to increase up to a height of 30 m (where it peaks at 3×10 4 cm 2 s −1 ), after which it decreases at a constant rate to a value of 2 cm May as there was a suite of other trace gas measurements made during this period (Lee et al., 2009), although it should be noted that there was little variability in the diurnal profiles of IO and BrO throughout the campaign (Fig. 1).The model is initialised at midnight and the evolution of iodine species, O 3 , NO x and HO x is followed for a period The predicted vertical profiles of the iodocarbons are shown in the supporting online text (Fig. S3, http://www.atmos-chem-phys-discuss.net/9/ 24281/2009/acpd-9-24281-2009-supplement.pdf).The longer lived iodocarbons (e.g.C 2 H 5 I) are better mixed through the BL, but the more photo-labile iodocarbons show a strong diurnal profile and are limited to <300 m (CH 2 I 2 , CH 2 IBr, CH 2 ICl).Using these measured iodocarbon fluxes, the predicted levels of daytime IO are only ∼30% of the observed values, peaking at only 0.45 ppt compared to the average observed value of around 1.5 ppt (Fig. 3).The predicted daytime O 3 depletion is only ∼2.9 ppb, compared to an observed decrease of ∼4.5 ppb observed at Cape Verde during May 2007 (Read et al., 2008).Additionally, the diurnal profile indicates that the IO should peak later in the evening (Fig. 3) due to a slow release of I atoms resulting from the relatively long lifetime of the iodocarbons (e.g.CH 2 IBr=∼15 min, CH 2 ICl=∼1 h, CH 3 I=23 h at midday) and hence slow build-up of IO during the daytime, in contrast to the observed top-hat profile (Figs. 1 and 3).These modelling results show that an additional flux of iodine is necessary in order to reproduce the observed levels of IO, as the iodocarbons can account for around one third of the observed IO mixing ratio.In the model, this additional source is considered to be in the form of an I 2 flux from the sea.Several laboratory studies have shown that deposition of O 3 on sea surface leads to the emission of I 2 (Garland and Curtis, 1981;Sakamoto et al., 2009).If a constant day and night I 2 flux of 5×10 7 molecules cm −2 s −1 is considered, the model predicts a peak of 1.6 ppt of IO during the day time, which is in agreement with the observations.In this scenario, the model also predicts a post-sunrise pulse of ∼1.4 ppt lasting for about 20 min.The best fit to the observed data is achieved by using a flux considered to Introduction

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Full Figure 4 shows the predicted diurnal vertical profiles for IO, O 3 , I 2 and OIO when this additional flux is included.Now, the levels of IO are in good accord with the observations at the height of measurements.The predicted vertical profile of IO shows a strong gradient with only 0.15 ppt (10%) seen at 500 m (Figs. 4 and 5).Even with the additional I 2 flux, the I 2 and OIO levels are under the detection limit throughout the BL.The K z profile in Fig. S2 considers that the K z reduces with height after 30 m but the model was found to be sensitive to the height from which the K z is considered to decrease.For an extreme condition, we assume that the K z does not decrease with height, but stays constant after 30 m.Using such a profile, the predicted [IO] at the height of measurements (10 m) is found to be ∼18% lower at the peak of the I atom flux due to faster mixing.Considering an O 3 deposition velocity of 0.05 cms −1 (Chang et al., 2004) and an O 3 mixing ratio of 35 ppb in the BL, the total flux of O 3 molecules into the ocean would be 4.3×10 10 molecules cm −2 s −1 .If the deposition of O 3 is responsible for the emission of I atoms from the ocean, this would mean that ∼3 iodine atoms are emitted to the gas phase for every 100 molecules of O 3 that are deposited on the ocean surface.One laboratory study has suggested that 20% of the O 3 deposited to the ocean surface reacts with I − ions in seawater to form molecular iodine and suggested that a significant fraction of this evaporates from the ocean surface (Garland and Curtis, 1980).The vertical profile of iodine species is shown in Fig. 5.It can be seen that the IO mixing ratio drops to about 10% of the value at the surface in the first 500 m, so that most of the IO is confined to the lower BL.The total column abundance predicted by the model is 9.7×10 11 molecules cm −2 , which is below the upper limit of 2×10 12 molecules cm −2 for current satellite measurements (Sch önhardt et al., 2007).

Effect of halogens on ozone
The vertical temperature profiles measured by the aircraft show a capped BL with a strong temperature inversion at about 1000 m during all the flights above Cape Verde indicating that the BL is decoupled from the free troposphere (Read et al., 2008).Read et al. also concluded that the MBL around Cape Verde is subject to minimal diurnal variations in dynamics because the temperature and wind speed are relatively constant.Furthermore, the free tropospheric O 3 mixing ratios over Cape Verde were found to be significantly lower than in the BL (Read et al., 2008), indicating that O 3 entrainment cannot take place into the BL directly above Cape Verde.O 3 entrainment into the BL actually takes place around 200 km further north-east of the measurement site, where measurements show elevated O 3 in the free troposphere (Lee et al., 2009); advection of air to Cape Verde then causes a rise in the O 3 concentrations during the night (Read et al., 2008).
During the summer month of May, the average loss of O 3 is ∼4.3 ppbd −1 , although in the absence of halogens, the model predicts a loss of only ∼2.5 ppb over the entire day at the height of measurements.This calculated O 3 loss is due to HO x chemistry and deposition to the ocean surface.If the observed levels of halogens are now included in the model (1.5 ppt IO and 2.5 ppt BrO), the total O 3 depletion at the height of measurements increases to ∼4.5 ppbd −1 (Fig. 3).The catalytic cycles involving halogens cause an increased loss of 1.6 ppbd −1 and the total predicted O 3 loss is now in good agreement with the observed loss. 1 km.This phenomenon is confirmed by the aircraft measurements, which showed that O 3 is well-mixed in the BL (Read et al., 2008).A point to note is the significant role that iodine chemistry plays in the O 3 depletion.At midday, the O 3 destruction rate due to iodine alone is 0.1 ppbh −1 , while the destruction due to bromine alone is 0.028 ppbh −1 .The combined impact of both halogens due to catalytic cycles involving IO and BrO (Gilles et al., 1997) destroys O 3 at a rate of 0.19 ppbh −1 .

Effect on HO x
In situ measurements of OH and HO 2 were performed at Cape Verde during the RHaM-BLe campaign using the LIF technique (Whalley et al., 2009).The measurements were made at the height of 10 m above sea level.Typical peak noon values of 0.25 ppt and 18 ppt were measured for OH and HO 2 , respectively.BrO and IO react with HO 2 to form HOBr and HOI, which on photolysis reproduce OH, thus reducing the HO x ratio (HO 2 /OH) ratio.Figure 6 shows the effect of halogen chemistry on the vertical profile of OH and HO 2 , for scenarios with and without halogens.At the height of the measurements, the model scenario without halogen chemistry over-predicts HO 2 by ∼14% of the measured concentration, and under-predicts the OH by ∼13%.Thus, the modelled HO x ratio is 94, compared with the measured ratio of 72 (Whalley et al., 2009).When halogen chemistry is included, the calculated OH and HO 2 levels and diurnal profile are in agreement with the observations: the modelled ratio is now 76. Figure 6 also shows that the vertical profile of OH and HO 2 are significantly different when halogens are included in the model, with the lowest predicted HO x ratio seen closer to the surface, although at the top of the BL the effect of halogens is greatly reduced.

DMS oxidation
The role of BrO in the oxidation of DMS has been discussed previously (e.g., Toumi, 1994;Saiz-Lopez et al., 2004;von Glasow and Crutzen, 2004;von Glasow et al., 2004).In the tropical MBL, the loss of DMS is thought to take place mainly through Introduction

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Full oxidation by OH (Turnipseed et al., 1996).However, BrO and to a lesser extent, IO can also cause the oxidation of DMS in the MBL.At the levels of BrO and IO observed at Cape Verde, the oxidation of DMS by BrO is comparable to OH but that by IO is much lower (Table 1).The oxidation of DMS by halogen oxides forms dimethyl sulfoxide (DMSO), which is either rapidly lost to background aerosols or forms methyl sulfinic acid (MSIA) rather than SO 2 and hence H 2 SO 4 .The presence of halogen oxides would therefore reduce the CCN formation potential of DMS if sulphuric acid formed from OH oxidation were assumed to build up to levels at which it would participate in nucleation (von Glasow and Crutzen, 2004).

Iodine oxide particle (IOP) formation
There are (at least) two major unresolved questions regarding IOPs: (i) the mechanism whereby iodine oxide molecules polymerize into sub-nm IOPs and (ii) the subsequent growth of these IOPs in the presence of condensable vapours (primarily water and acids) in the marine atmosphere.In the past, the oxidation of iodine oxides by O 3 has been suggested to lead to the formation of I 2 O 4 or I 2 O 5 , which can then polymerise to form IOPs (Jimenez et al., 2003;Saunders and Plane, 2005).However, very recent experiments at the University of Leeds show that IOPs are also produced from IO in the absence of O 3 (Saunders et al., 2009).This indicates that the particles form from the polymerization of I 2 O 3 and I 2 O 4 , which are produced from the recombination of IO and OIO or the self-association of OIO, respectively (J. C. G ómez Martin, University of Leeds, personal communication, 2009).Quantum calculations, using the same basis set for atomic I described in Kaltsoyannis and Plane (2008), show that I 2 O 3 does not dimerize effectively with itself, but forms a much more stable complex with I 2 O 4 .I 2 O 4 also forms a stable dimer with itself.Hence, I 2 O 4 appears to be the key species for initiating the formation of IOPs in the MBL.In THAMO, I 2 O 4 is treated as the primary polymerising species, on which I 2 O 3 can condense.Both species can condense on larger polymers, which can all also coagulate with each other.These reactions are assumed to occur at their respective collision frequencies (or kernels), which is neces-Introduction

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Interactive Discussion sary to model the explosive growth of IOPs observed in the laboratory (Saunders and Plane, 2006).The total number of particles growing to diameter=2 nm are calculated using a semiimplicit, volume conserving coagulation routine as described by Jacobson (2005).Along with growth due to coagulation, I 2 O 3 , I 2 O 4 and IOPs are also lost to background aerosols, the rate of which is calculated according to the molecular transfer approximation (Fuchs, 1964).The THAMO model predicts that I 2 O 4 formation is important only during the night at the IO mixing ratios observed at Cape Verde.This is because I 2 O 4 forms from the dimerisation of OIO, and OIO photolyzes very rapidly as can be seen in Fig. 4 (G ómez Martin et al., 2009).In the MBL, IOPs can potentially grow much faster due to the uptake of acids and water vapour and there is no reason to expect that condensable vapours in the MBL would not condense onto freshly nucleated particles.Sulphuric acid concentrations in the remote MBL range between 0.01-0.03ppt (e.g., Berresheim et al., 1993), and in the presence of sufficient ammonia (NH 3 ), will condense onto any particles that are formed.Indeed recent laboratory experiments at the University of Leeds show that H 2 SO 4 is taken up on humidified IOPs with a mass accommodation coefficient of 0.75 at 90% relative humidity (Saunders et al., 2009).The uptake of H 2 SO 4 on ultrafine particles occurs once the particles are large enough to become "activated"; past studies on other atmospheric particles have shown that the activation size ranges between 1.5 and 3 nm (Kulmala et al., 2007).For simplicity, we do not consider a diurnal profile in the [H 2 SO 4 ] and assume that uptake of ammonia (NH 3 ) will keep the pH of the IOPs close to neutral, so that further uptake of H 2 SO 4 is not inhibited.This assumption is valid as the background NH 3 mixing ratio in the MBL ranges between 25 and 75 ppt (Quinn et al., 1988), and aliphatic amines have been observed in sub-µm particles around Cape Verde (M üller et al., 2009).
Results from the THAMO runs using this IOP formation and growth mechanism are illustrated in Fig. 7.The figure shows the number of particles at the measurement height of 10 m, which might grow large enough (taken to be >20 nm) to have a good chance of surviving to become CCN.These are the number of potential CCN formed Introduction

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Full at 10 m after 24 h, as a function of the mean daytime IO concentration at 10 m.The IO is changed by altering the sea-to-air flux of I 2 species in addition to the halocarbon flux as mentioned in Sect.4.2.Two cases are illustrated in the figure to show the sensitivity to the background aerosol surface area (ASA), which is set to a typical remote MBL value of 1×10 −6 cm 2 cm −3 or an ultra-clean value of 1×10 −7 cm 2 cm −3 .
The figure shows rather graphically that the predicted number of particles is highly sensitive to (i) the background ASA, because the loss to background aerosols through uptake is usually faster than growth by coagulation and condensation (except at high [IO]); and (ii) the IO concentration because the rates of formation of OIO and I 2 O 4 depend non-linearly on [IO].The figure also demonstrates the probability of forming potential CCN at Cape Verde is vanishingly small.The model predicts a maximum of only ∼1×10 −25 particles cm −3 (diameter=20 nm) at the height of measurements when the ASA is 1×10 −6 cm 2 cm −3 , and ∼1×10 −17 when the ASA is 2×10 −7 cm 2 cm −3 .Thus, at the concentrations of IO seen at Cape Verde, the formation of CCN from iodine chemistry is extremely unlikely.However, the model results show that for a higher [IO], formation of IOPs could be an important source of new particles in the MBL.Such bursts of new particles (diameter=3-10 nm) have been observed in coastal areas peaking at around 2.5×10 5 particles cm −3 (e.g., Whitehead et al., 2009).In these regions [IO] can reach values as high as ∼27 ppt (e.g., Whalley et al., 2007).Similar bursts have also been observed in Antarctica, where >4000 particlescm −3 were observed in the summer (Davison et al., 1996), when the [IO] peaks about 10 ppt (Saiz-Lopez et al., 2007).
OIO), as well as the night-time radical NO 3 .All three species were below the detection limit of the instrument throughout the campaign i.e. [I 2 ]≤15 ppt, [OIO]≤10 ppt, and [NO 3 ]≤4 ppt.
has suggested that deposition of O 3 on the ocean surface could lead to the emission of iodocarbons in the presence of dissolved organic matter, which on photolysis could contribute to RIS in the MBL.Fluxes of iodocarbons measured over the open ocean close to the measurement site, as a part of the RHaMBLe campaign, confirm their presence(Jones et al., 2009)  and indicate that they could indeed act as a source of RIS in the remote MBL.
Furthermore, the best fit of the I 2 flux to the measured IO diurnal profile requires that the emission of I atoms from the ocean surface is more active during the daytime (Fig. S4, http://www.atmos-chem-phys-discuss.net/9/24281/2009/ acpd-9-24281-2009-supplement.pdf), which could indicate a photochemical production of I 2 or other iodocarbons in the sea-surface micro layer, perhaps coupled with the Figure 3 also shows that this observed depletion is well mixed in the BL with a maximum difference of 1.2 ppb between the surface and at

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Summary and conclusionsMeasurements of halogen species at an open ocean site in the tropical Atlantic Ocean are presented.IO and BrO are ubiquitously detected at Cape Verde in the presence of sunlight with no apparent dependence of wind speed or direction.Both radicals display a distinct top-hat profile with the maximum values seen between 09:00-17:00 GMT and

Fig. 1 .Fig. 4 .Fig. 5 .Fig. 6 .Fig. 7 .
Fig. 1.The monthly averaged diurnal profiles for IO and BrO are shown in panels (a) and (b).Both radicals show a distinctive top-hat profile with the highest mixing ratios measured between 09:00-17:00 GMT.The average annual profile of IO and BrO along with the highest measured mixing ratio for each month is shown in panel (c) indicating little annual variability.
., Introduction Initially, the injection of iodine into the model is constrained using the average sea-air iodocarbon fluxes measured over the open ocean close to Cape Verde(Jones et al., 2009): CH 2

Table 1 .
annual profile of both the species shows little variation.The presence of IO and BrO above the detection limit of the DOAS instrument indicates that halogen chemistry could play a wider role in the open ocean MBL.Using the 1-D model THAMO, it is shown that the observed iodocarbon fluxes cannot explain the IO measurements, suggesting that other iodine compounds must be emitted from the ocean surface.The modelling results also show that the observed O 3 depletion cannot be explained in the absence of halogen chemistry.The model is able to account for the low HO x ratio measured at Cape Verde, indicating that halogen chemistry directly affects the oxidising capacity of the atmosphere.Bromine chemistry is shown to oxidise DMS at rates comparable to OH affecting the CCN formation potential of DMS.Finally, modelling results suggest that IOP formation does not occur at Cape Verde and higher IO concentrations are necessary to produce particles of a climatically important size.However, further studies are necessary in order to fully understand the potential of iodine chemistry to form CCN in other open ocean environments.Introduction Rates of DMS oxidation calculated for observations at Cape Verde. the