ACPAtmospheric Chemistry and PhysicsACPAtmos. Chem. Phys.1680-7324Copernicus PublicationsGöttingen, Germany10.5194/acp-16-15593-2016Nighttime atmospheric chemistry of iodineSaiz-LopezAlfonsoa.saiz@csic.eshttps://orcid.org/0000-0002-0060-1581PlaneJohn M. C.https://orcid.org/0000-0003-3648-6893CuevasCarlos A.MahajanAnoop S.https://orcid.org/0000-0002-2909-5432LamarqueJean-Françoishttps://orcid.org/0000-0002-4225-5074KinnisonDouglas E.Department of Atmospheric Chemistry and Climate, Institute of Physical
Chemistry Rocasolano, CSIC, Madrid, SpainSchool of Chemistry, University of Leeds, Leeds, UKIndian Institute of Tropical Meteorology, Pune, IndiaAtmospheric Chemistry Observations & Modeling Laboratory, National Center for Atmospheric Research, Boulder, 10 Colorado, USAAlfonso Saiz-Lopez (a.saiz@csic.es)19December20161624155931560418May201615June20162December20165December2016This work is licensed under a Creative Commons Attribution 3.0 Unported License. To view a copy of this license, visit http://creativecommons.org/licenses/by/3.0/This article is available from https://acp.copernicus.org/articles/16/15593/2016/acp-16-15593-2016.htmlThe full text article is available as a PDF file from https://acp.copernicus.org/articles/16/15593/2016/acp-16-15593-2016.pdf
Little attention has so far been paid to the nighttime atmospheric chemistry
of iodine species. Current atmospheric models predict a buildup of HOI and
I2 during the night that leads to a spike of IO at sunrise, which is not
observed by measurements. In this work, electronic structure calculations are
used to survey possible reactions that HOI and I2 could undergo at night
in the lower troposphere, and hence reduce their nighttime accumulation. The
new reaction NO3+ HOI → IO + HNO3 is proposed, with a
rate coefficient calculated from statistical rate theory over the temperature
range 260–300 K and at a pressure of 1000 hPa to be k(T)
= 2.7 × 10-12
(300 K/T)2.66 cm3 molecule-1 s-1. This reaction is
included in two atmospheric models, along with the known reaction between
I2 and NO3, to explore a new nocturnal iodine radical activation
mechanism. The results show that this iodine scheme leads to a considerable
reduction of nighttime HOI and I2, which results in the enhancement of
more than 25 % of nighttime ocean emissions of HOI + I2 and the
removal of the anomalous spike of IO at sunrise. We suggest that active
nighttime iodine can also have a considerable, so far unrecognized, impact on
the reduction of the NO3 radical levels in the marine boundary layer
(MBL) and hence upon the nocturnal oxidizing capacity of the marine
atmosphere. The effect of this is exemplified by the indirect effect on
dimethyl sulfide (DMS) oxidation.
Introduction
Active nighttime iodine chemistry was first evidenced a decade ago when it
was shown that nocturnal I2 emitted by macroalgae could react with
NO3 leading to the formation of IO and OIO, which were measured in the
coastal marine boundary layer (MBL) at Mace Head, Ireland (Saiz-Lopez and Plane, 2004). The nitrate
radical has also been recently suggested as a nocturnal loss of
CH2I2, which helps to reconcile observed and modelled
concentrations of this iodocarbon over the remote MBL (Carpenter et al.,
2015). However, most of the work on reactive atmospheric iodine has focused
on the use of daytime observations and models to assess its role in the
catalytic destruction of ozone and the oxidizing capacity of the troposphere
(e.g. Saiz-Lopez et al., 2012b, and references therein). In the MBL,
iodine-catalysed along with bromine-catalysed ozone destruction contributes up to
45 % of the observed daytime depletion (Read et al., 2008; Mahajan et
al., 2010a), although this contribution shows large geographical variability
(Mahajan et al., 2012; Gómez Martín et al., 2013; Prados-Roman et
al., 2015b; Volkamer et al., 2015). Iodine compounds have also been
implicated in the formation of aerosols, although the mechanisms and
magnitudes of these processes are not fully understood (Hoffmann et al.,
2001; O'Dowd et al., 2002; McFiggans et al., 2004; Gómez Martín et
al., 2013; Allan et al., 2015; Roscoe et al., 2015). Reactive forms of
inorganic iodine may also contribute to the oxidation of elemental mercury
over the tropical oceans (Wang et al., 2014). In recent years, iodine sources
and chemistry have also been implemented in global models demonstrating the
effect of iodine chemistry in the oxidation capacity of the global marine
troposphere
(Ordóñez et al., 2012; Saiz-Lopez et al., 2012a,
2014; Sherwen et al., 2016).
Iodine is emitted into the atmosphere from the ocean surface in both organic
and inorganic forms. The main organic compounds emitted are methyl iodide
(CH3I), ethyl iodide (C2H5I), propyl iodide (1- and
2-C3H7I), chloroiodomethane (CH2ICl), bromoiodomethane
(CH2IBr) and diiodomethane (CH2I2) (Carpenter, 2003;
Butler et al., 2007; Jones et al., 2010; Mahajan et al., 2012). However,
these organic compounds contribute only up to a fourth of the MBL iodine
loading (Jones et al., 2010; Mahajan et al., 2010a; Großmann et al.,
2013; Prados-Roman et al., 2015b). Inorganic emissions of HOI and I2,
which result from the deposition of O3 at the ocean surface and
subsequent reaction with I- ions in the surface microlayer, account for
the main source of iodine in the MBL (Carpenter et al., 2013).
Recent laboratory experiments have shown that HOI is the major compound
emitted and provided parameterizations of the fluxes of both species
depending on wind speed, temperature, and the concentrations of O3 and
I- (Carpenter et al., 2013; MacDonald et al., 2014). These
parameterized fluxes of HOI and I2 have then been used in a
one-dimensional model to study the diurnal evolution of the IO and I2
mixing ratios at the Cape Verde Atmospheric Observatory (CVAO) (Carpenter
et al., 2013; Lawler et al., 2014). The model simulations replicate well the
levels and general diurnal profiles of IO and I2, although an early
morning “dawn spike” in IO is predicted by the models but has not been
observed (Read et al., 2008; Mahajan et al., 2010a). The morning peak
predicted by current iodine chemistry models is due to a buildup of the
emitted I2 and HOI (which is converted into IBr / ICl through
heterogeneous sea-salt recycling) over the course of the night, followed by
rapid photolysis at sunrise.
Traditionally it has been thought that iodine chemistry has a negligible
effect on oxidizing capacity of the nocturnal marine atmosphere. As a
consequence, unlike the demonstrated effect of iodine on the levels of
daytime oxidants, the impact of active iodine upon the main nighttime
oxidant, NO3, remains an open question. This is important given that in
many parts of the ocean the NO3+ dimethyl sulfide (DMS) reaction is at least as important
as OH + DMS in oxidizing DMS (Allan et al., 2000), and hence a reduction of
NO3 may have an effect in the production of SO2 and methane
sulfonic acid (MSA). Here, we discuss possible mechanisms of nighttime iodine
radical activation and their potential effect on nighttime iodine ocean
fluxes and the currently modelled dawn spike in IO. A new reaction of HOI with
NO3 is proposed, supported by theoretical calculations. We explore the
implications of this new reaction both for iodine and NO3 chemistries.
Nocturnal iodine radical activation mechanism
We use the reaction mechanism that has recently been described in the global
modelling studies by Saiz-Lopez et al. (2014) and Ordóñez et al. (2012)
(see Supplement). In addition to the reactions included
in that scheme, we also include nighttime gas-phase reactions based on the
theoretical calculations described below. The additional reactions are listed
in Table 1, and a scheme with this new nocturnal chemistry is included in
Fig. 1.
New nocturnal iodine chemistry (in white) implemented in the THAMO
and CAM-Chem models.
Nighttime reactions of emitted inorganic iodine compounds
considered in addition to the iodine chemistry scheme used by
Saiz-Lopez et al. (2014).
No.ReactionNotesReaction (R1)I2+ NO3→ I + IONO21.5 × 10-12 cm3 molecule-1 s-1 (Chambers et al., 1992).Reaction (R2)HOI + NO2→ I + HNO3Endothermic by 9 kJ mol-1 and the transition state is 73 kJ mol-1 above the reactants.Reaction (R3)HOI + HNO3→ IONO2+ H2OExothermic by 11 kJ mol-1. The reaction first forms a complex 21 kJ mol-1 below the reactants, but this rearranges to the products via a transition state that is 110 kJ mol-1 above the reactants.Reaction (R4)HOI + NO3→ IO + HNO3Exothermic by 11 kJ mol-1 with all transition states below the reactants. k(T)= 2.7 × 10-12 (300 K/T)2.66 cm3 molecule-1 s-1.
To the best of our knowledge, reactions of HOI specific to nighttime have
not been studied, either theoretically or through laboratory experiments.
Currently, HOI is thought to build up overnight until sunrise, with only
heterogeneous uptake on sea-salt aerosol as a nighttime loss process
(Saiz-Lopez et al., 2012b; Simpson et al., 2015). In addition to the well
known
I2+NO3→I+IONO2
(Chambers et al., 1992), here we consider several possible HOI reactions that
could occur at night, in the absence of photolysis and OH:
HOI+NO2→I+HNO3HOI+HNO3→IONO2+H2OHOI+NO3→IO+HNO3.
Theoretical calculations
In order to explore the feasibility of Reactions (R2)–(R4) taking place under the
conditions of the lower troposphere, we carried out electronic structure
calculations using the hybrid density functional/Hartree–Fock B3LYP method
from within the Gaussian 09 suite of programs (Frisch et al., 2009),
combined with a G2 level basis set for I (Glukhovtsev et al., 1995) and the
standard 6–311 + g(2d,p) triple-zeta basis set for O, N and H. Following
geometry optimizations of the relevant points on the potential energy
surfaces, and the determination of their corresponding vibrational
frequencies and (harmonic) zero-point energies, energies relative to the
reactants were obtained at the same level of theory. Spin–orbit corrections
of -30.0 (Mečiarová et al., 2011), -14.4 (Khanniche et al., 2016),
-5.9 (Šulková et al., 2013) and -4.8 (Kaltsoyannis and Plane, 2008)
kJ mol-1 were applied to the energies of I, IO, HOI and IONO2,
respectively.
Reaction (R2) is endothermic by 2.6 kJ mol-1, and therefore within the
expected error of ±10 kJ mol-1 at this level of theory. Hence, it might be
reasonably fast. However, the transition state of the reaction, which is
illustrated in Fig. 2a, is 73 kJ mol-1 above the reactants, so this
reaction will not occur at tropospheric temperatures. Reaction (R3) is
exothermic by 19.8 kJ mol-1. An HOI–HNO3 complex first forms
(Fig. 2b), which is 21 kJ mol-1 below the reactants. However, this
complex re-arranges to the IONO2+ H2O products via the cyclic
transition state shown in Fig. 2c, which is 110 kJ mol-1 above the
reactants.
(a) Transition state for the reaction between HOI and NO2 to
form HNO3+ I and (b) complex formed between HOI and HNO3, which
then reacts via transition state (c) to form IONO2+ H2O.
The stationary points on the potential energy surface (PES) for Reaction (R4)
are illustrated in Fig. 3. HOI and NO3 associate to form a complex which
is 24 kJ mol-1 below the reactant entrance channel. Hydrogen atom transfer
involves a submerged transition state to form an IO–HNO3 complex, which
can then dissociate to the products IO + HNO3. The vibrational
frequencies, rotational energies and geometries (in Cartesian co-ordinates)
of these intermediates are listed in Table 2. Overall, the reaction is
exothermic by 14 kJ mol-1. The energies of the HOI–NO3 complex
and the transition state are assigned the same spin–orbit correction as HOI
(-5.9 kJ mol-1; Šulková et al., 2013), whereas the
IO–HNO3 complex is assigned the spin–orbit correction of IO
(-14.4 kJ mol-1; Khanniche et al., 2016). This reflects the H–OI
bond only increasing from 0.97 Å in HOI to 1.1 Å in the transition
state, compared with 1.7 Å in the IO–HNO3 complex. The spin–orbit
correction for the transition state is therefore likely to be closer to that
of HOI. Assigning the HOI spin–orbit correction therefore means that the
barrier is highest with respect to the reactants so that the estimated rate
coefficient (see below) may be a lower limit.
Calculated vibrational frequencies, rotational constants and
energies of the stationary points and asymptotes on the HOI + NO3
doublet potential energy surface.
SpeciesGeometryaVibrational frequenciesbRotational constantscPotentialenergydHOI + NO3603, 1084, 3803 and 261, 261, 805,1108, 1108, 1126623.9, 8.182, 8.076 and 13.84, 13.84, 6.9190.0IOH–NO3 complexO 1.623, 0.284, -0.331 H 1.484, -0.657, -0.043 I 0.009, 1.205, 0.286 N -0.456, -2.265, 0.030 O -1.052, -3.321, -0.0473 O -1.147, -1.195, -0.228 O 0.742, -2.161, 0.333 55, 84, 118, 161, 196, 615, 629, 667, 705, 803, 968, 1228, 1273,1491, 32685.610, 0.916, 0.806-24.0IO–H–NO2 TSO 0.309, 1.515, 0.247 H -0.834, 1.314, -0.017 I 1.280, -0.089, -0.093 N -2.349, -0.133, 0.019 O -3.518, -0.429, -0.035 O -1.444, -0.962, 0.257 O -2.019, 1.117, -0.187 1249i, 70, 97, 103, 225, 472, 676, 698, 797, 806, 1041, 1147, 1308, 1513, 16266.300, 0.864, 0.767-16.4IO–HNO3 complexO 0.571, 1.350, 0.348 H -1.111, 1.098, -0.020 I 1.870, 0.0645, -0.152 N -2.503, -0.202, 0.0186 O -3.673, -0.396, -0.170 O -1.654, -0.986, 0.401 O -2.081, 1.090, -0.242 35, 43, 76, 126, 198, 623, 677, 703, 772, 798, 939, 1331, 1416, 1713, 32817.058, 0.605, 0.566-34.8IO + HNO3648 and 477, 585, 649, 782, 901, 1320, 1345, 1738, 37249.844 and 13.01, 12.05, 6.258-10.6
a Cartesian co-ordinates in Å. b Given as wavenumber in
cm-1. c In GHz. d In kJ mol-1, including zero-point
energy and spin–orbit coupling of I and IO (see text).
The rate coefficient for Reaction (R4) was then estimated using
Rice–Ramsperger–Kassel–Markus (RRKM) theory, employing a multi-well
energy-grained master equation solver based on the inverse Laplace transform
method – MESMER (Master Equation Solver for Multi-well Energy Reactions)
(Roberston et al., 2014). The reaction proceeds
via the formation of the excited HOI–NO3 complex from HOI + NO3.
This complex can then dissociate back to the reactants or
rearrange to the IO–HNO3 intermediate complex over the transition
state, which can in turn dissociate to the products IO + HNO3. Either
of the intermediates can also be stabilized by collision with the third body
(N2). The time evolution of all these possible outcomes is modelled
using the master equation.
The internal energies of the intermediates on the PES were divided into a
contiguous set of grains (width 10 cm1), each containing a bundle
of rovibrational states calculated with the molecular parameters in Table 2,
using the rigid-rotor harmonic oscillator approximation for all species. It
should be noted that the HOI–NO3 and IO–HNO3 complexes both have
low-frequency vibrational modes (< 100 cm-1) which should more
correctly be treated as hindered rotors rather than vibrations. However, in
our experience this is not worth doing this until experimental rate
coefficients are available to fit the rotor barrier heights. In any case,
the energies of both complexes are far enough below the energy of the
entrance channel (Fig. 3) that relatively small changes in their densities
of states will have a minor effect on the overall rate coefficient. Each
grain was then assigned a set of microcanonical rate coefficients linking it
to other intermediates, calculated by RRKM theory. For dissociation to
products or reactants, microcanonical rate coefficients were determined
using inverse Laplace transformation to link them directly to the capture
rate coefficient, kcapture. For Reaction (R4) and the reverse
reaction IO + HNO3 involving neutral species, kcapture
was set to a typical capture rate coefficient of 2.5 × 10-10
(T/300 K)1/6 cm3 molecule-1 s-1, where the
small positive temperature dependence is characteristic of a long-range
potential governed by dispersion and dipole–dipole forces
(Georgievskii and Klippenstein, 2005).
Potential energy surface for the reaction between HOI and
NO3, which contains two intermediate complexes separated by a submerged
barrier.
The probability of collisional transfer between grains was estimated using
the exponential down model, where the average energy for downward transitions
was set to < ΔE>down= 300 cm-1
for N2 as the third body (Gilbert and Smith, 1990). MESMER
determines the temperature- and pressure-dependent rate coefficient from the
full microcanonical description of the system time evolution by performing an
eigenvector/eigenvalue analysis (Bartis and Widom, 1974). The resulting rate
coefficient over the temperature range 260–300 K at a pressure of 1000 hPa
is k4(T)= 2.7 × 10-12 (300 K/T)2.66 cm3 molecule-1 s-1. Because the
intermediate complexes are not strongly bound, and the transition state and
products are below the entrance channel, the only products formed in Reaction (R4)
under atmospheric conditions are IO + HNO3. The uncertainty in
k4 arises principally from the estimated capture rate
coefficient (see above) and the height of the barrier below the entrance
channel. As discussed above, the spin–orbit correction of the transition
state is likely to be larger than the value of -5.9 kJ mol-1
corresponding to HOI, so k4 is possibly a lower limit. For
instance, if the barrier height is decreased by 3 kJ mol-1,
k4 increases by a factor of 1.9. If the barrier is lower by
8.5 kJ mol-1 (corresponding to the transition state having the same
spin–orbit correction as IO), then k4 would increase by a
factor of 5.1. Nevertheless, noting that the capture rate coefficient could
be lower – perhaps by a factor of 2 – than the estimate used here, we prefer
to use the value for k4 calculated with the potential surface
in Fig. 3. Of course, if k4 is larger, then the atmospheric
impacts of Reaction (R4) discussed in Sect. 4 will be even more pronounced.
Note that NO3 also reacts with CH2I2 with a rate constant
∼ 2–4 × 10-13 cm3 molecule-1 s-1, which
can have a significant effect on nighttime CH2I2 concentration
(Carpenter et al., 2015). However the products of this reaction are still
uncertain (Nakano et al., 2006; Carpenter et al., 2015) and its rate is
considerably slower than that of Reaction (R4).
In summary, the only likely gas-phase reactions that I2 and HOI undergo
in the nighttime troposphere are Reactions (R1) and (R4), respectively. These
are included in the model reaction scheme to examine their impacts on the
evolution of iodine species in the atmosphere.
Atmospheric modelling
We use two atmospheric chemical transport models to study (i) the
impact of this new chemistry on the nighttime chemistry and partitioning of
iodine species and (ii) the resulting geographical distribution of
nocturnal iodine and impact on NO3 within the global marine boundary
layer.
The first model, Tropospheric HAlogen chemistry MOdel (THAMO) (Saiz-Lopez et al., 2008), is used for a
detailed kinetics study of the impact of the different reactions shown in
Table 1 as well as to assess which uptake rates best reproduce observations
from a field study at the CVAO (Carpenter et al., 2011). THAMO has been used
in the past to study iodine chemistry at the CVAO, and further details
including the full chemical scheme can be found elsewhere (Saiz-Lopez et al., 2008; Mahajan et al., 2009, 2010a, b; Lawler et al.,
2014). Briefly, THAMO 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, O3, NOx and HOx chemistry and is constrained
with typical measured values of other chemical species in the MBL: [CO] =
110 nmol mol-1; [DMS] = 30 pmol mol-1; [CH4] =
1820 nmol mol-1; [ethane] = 925 pmol mol-1; [CH3CHO] =
970 pmol mol-1; [HCHO] = 500 pmol mol-1; [isoprene] =
10 pmol mol-1; [propane] = 60 pmol mol-1; [propene] =
20 pmol mol-1. The average background aerosol surface area (ASA) used
is 1 × 10-6 cm2 cm-3 (Read et al., 2008, 2009; Lee
et al., 2009, 2010). The model is initialized at midnight and the evolution
of iodine species, O3, NOx and HOx is followed until the model
reaches steady state.
THAMO modelled diurnal variation of HOI, I2 (upper panels) and
the HOI / I2 flux from the ocean surface (bottom panel). The right-hand
panels are from Scenario 1, which do not include nighttime reactions of HOI
and I2 with NO3, while the left-hand panels include the reactions
in Scenario 2. In bottom panel red lines represent Scenario 1, while black
lines correspond to Scenario 2.
The second model is the global 3-D chemistry-climate model CAM-Chem (Community
Atmospheric Model with chemistry, version 4.0), which is used to study the
impact of Reaction (R1) and (R4) on a global scale. The model includes a
comprehensive chemistry scheme to simulate the evolution of trace gases and
aerosols in the troposphere and the stratosphere (Lamarque et al., 2012). The
model runs with the iodine, bromine and chlorine chemistry schemes from previous
studies (Fernandez et al., 2014; Saiz-Lopez et al., 2014,
2015), including the photochemical breakdown of bromo- and iodocarbons
emitted from the oceans (Ordóñez et al., 2012) and abiotic oceanic
sources of HOI and I2 (Prados-Roman et al., 2015a). CAM-Chem has been
configured in this work with a horizontal resolution of 1.9∘
latitude by 2.5∘ longitude and 26 vertical levels, from the
surface to ∼ 40 km altitude. All model runs in this study were
performed in the specified dynamics mode (Lamarque et al., 2012) using
offline meteorological fields instead of an online calculation, to allow
direct comparisons between different simulations. This offline meteorology
consists of a high-frequency meteorological input from a previous free
running climatic simulation.
It should be noted that during nighttime the uptake on aerosols of emitted
species such as I2 and HOI, and the uptake of reservoir species such as
IONO2, can play a major role in the cycling of iodine. Observations at
CVAO show that I2 peaked at about 1 pmol mol-1 during the night and that
ICl was not detected above the 1 pmol mol-1 detection limit of the instrument
(Lawler et al., 2014). In order to match these
observations, we need to reduce the uptake and heterogeneous recycling of
iodine species. The uptake rates of chemical species on the background
sea-salt aerosols are determined by their uptake coefficients
(γ). The database of mass accommodation and/or uptake
coefficients is rather sparse and essentially limited to I2, HI, HOI,
ICI and IBr on pure water/ice and on sulfuric acid particles
(Sander et al., 2006). Other iodine species
which are likely to undergo uptake onto aerosol are OIO, HIO3,
INO2, IONO2 and I2O2 (Saiz-Lopez et al., 2012a;
Sommariva et al., 2012). Uptake of HOI is very uncertain, with
γ(HOI) ranging from 2 × 10-3 to 0.3 depending
on the surface composition and state (Holmes et al.,
2001). Sommariva et al. (2012) assumed γ(HOI) to be 0.6, similar to the value for HOBr measured by
Wachsmuth et al. (2002). In the case of IONO2, the
uptake coefficient has not been measured, with most models using values of
0.1 (von Glasow et al., 2002; Saiz-Lopez et al., 2008; Mahajan et al.,
2009, 2010a, b; Leigh et al., 2010; Sommariva et al., 2012; Lawler et al., 2014). The modelled levels of I2
and ICl change with different values of uptake coefficients. To match the
CVAO I2 and ICl observations (Lawler et al., 2014), we have used
γ= 0.01 for HOI and IONO2, which is within the
uncertainty in the literature, and assumed that 80 % is recycled as
I2. Further measurements of these dihalogen species are needed to better
constrain their heterogeneous recycling on sea-salt aerosols.
THAMO modelled diurnal variation of IO, NO3 and the
IONO2. The right-hand panels are from Scenario 1, which do not include
nighttime reactions of HOI and I2 with NO3, while the left-hand
panels include the reactions in Scenario 2.
Sensitivity run showing the effect of the uncertainty in the rate
constant estimation on the reduction of NO3 peak nighttime
concentration at the surface – the red point is the theoretical estimate.
Results and discussion
Of the possible nocturnal iodine activation reactions involving the inorganic
iodine source gases I2 and HOI, only Reactions (R1) and (R4) appear to
be likely candidates (see Sect. 3). We therefore designed two modelling
scenarios: Scenario 1 (S1), without nighttime reactions of I2 or HOI
with NO3; and Scenario 2 (S2), including Reactions (R1) and (R4) for the
degradation of HOI and I2 by NO3. In the one-dimensional model
THAMO, the I2 and HOI are injected into the atmosphere from the ocean
surface using the flux parameterizations derived from laboratory experiments
(Carpenter et al., 2013; MacDonald et al., 2014). Figure 4 shows the
resulting diurnal evolution of the HOI and I2 mixing ratios in the two
scenarios, after 2 days of simulation time. The I2 mixing ratio peaks
during the night in both the scenarios due to quick loss by photolysis during
the daytime. By contrast, HOI is present during daytime due to its production
through the reaction of IO with HO2 and peaks just before sunset. In
the first scenario, without the inclusion of Reactions (R1) and (R4),
Fig. 4 (panels on right-hand side) shows that I2 builds up during the
night, reaching a concentration peak just before dawn. This is especially
noticeable as the daytime concentrations are much lower than during the
night. On the other hand, HOI concentrations decrease during night until
dawn, when they drop to zero. For both species, inclusion of reactions with
NO3 causes a decrease in their respective nocturnal concentrations
(Fig. 4, panels on left-hand side). The inclusion of Reactions (R1) and (R4)
also leads to a modelled I2 concentration which is in better agreement
with the observations of the molecule made at CVAO (Lawler et al., 2014),
reaching peak values of about 1 pmol mol-1, as compared to about
3 pmol mol-1 for the scenario without nighttime reactions. An
additional consequence of including Reactions (R1) and (R4) is the
significant increase of the sea–air fluxes of HOI and I2 at night due to
their atmospheric removal by NO3 (Fig. 4, bottom panel).
Modelled annual average of HOI (a) and I2(b) during nighttime (from 00:00 to 01:00 LT) at the surface level. The panels show the
difference in volume mixing ratio between the simulations with and without
Reactions (R1) and (R4).
Modelled annual average of IONO2(a) and NO3(b) during
nighttime (from 00:00 to 01:00 LT) at the surface level, as the difference in
volume mixing ratio between the simulations with and without Reactions (R1)
and (R4).
Increase in the DMS levels during nighttime (from 00:00 to 01:00 LT)
at the surface level due to the inclusion of the Reactions (R1) and (R4) in
CAM-Chem.
Figure 5 shows the diurnal evolution of IO, NO3 and IONO2 in both
model scenarios after 2 days of simulation time. Although the daytime peak
values of IO are well reproduced in both scenarios, reaching about
1.5 pmol mol-1 around noon similar to the ground-based observations
(Read et al., 2008), the inclusion of Reactions (R1) and (R4) leads to the
removal of the dawn spike in IO, which is predicted by current iodine models
but was not observed at CVAO (Read et al., 2008; Mahajan et al., 2010a). The
IO dawn spike predicted by models is due to a buildup of the emitted I2
and HOI (which is converted into IBr / ICl through heterogeneous
recycling) over the night, followed by rapid photolysis after first sunlight.
However, due to the considerable removal of HOI and I2 through the night
due to reaction with ambient NO3, this spike does not appear in the
second scenario, leading to a modification of the diurnal profile of IO that
better matches with observations.
Hourly averaged concentration of HOI, IONO2 and I2 in
the Mediterranean Sea at the surface level (long: 10 → 20∘ E, lat: 33 → 40∘ N)
Hourly averaged concentration of HOI, IONO2 and I2
(upper panel) and NO3 (bottom panel) in the Pacific Ocean at the south
of Baja California Peninsula at the surface level (long:
-110 →-106∘ E,
lat: 16 → 23∘ N)
Reactions (R1) and (R4) also reduce the NO3 mixing ratio (Fig. 5, middle
panels). In Scenario 1, the NO3 is modelled to peak at about 14 pmol mol-1
just before dawn. However, the inclusion of Reactions (R1) and (R4) leads to near
complete depletion of NO3 close to the surface, with the peak level at
the surface reaching only 2 pmol mol-1, since Reactions (R1) and (R4) become the
main atmospheric loss processes for NO3 in the lower MBL. These
reactions lead however to the buildup of IONO2 during the night (Fig. 5,
bottom panels). In the absence of Reactions (R1) and (R4), significant levels of
IONO2 are seen only at dawn and dusk since no other reactions produce
IONO2 at night, and during the day IONO2 is removed by photolysis.
However, with continuous conversion of I2 and HOI to IONO2 by
Reactions (R1) and (R4) in Scenario 2, IONO2 is modelled to reach up
to 3 pmol mol-1 in the nocturnal MBL.
Given the associated uncertainty in the theoretical estimate of the
k4, we used THAMO to assess the sensitivity of surface NO3
to k4. Figure 6 shows that NO3 peak nighttime
concentration is in fact highly coupled to k4, with the
expected uncertainty in k4 of 1 order of magnitude (see
above) giving rise to a factor-of-2 change in NO3. A laboratory
measurement of k4 should therefore be undertaken in the future.
We now implement the nighttime reactions in the 3-D global model (CAM-Chem)
to assess the resulting geographical distributions and impacts of these
reactions. We have also run two different scenarios in CAM-Chem, the first
without Reactions (R1) and (R4) in the chemical scheme, and the second
including the new nighttime iodine chemistry. Figure 7 shows how the
inclusion of Reactions (R1) and (R4) reduces globally the nighttime
concentrations of I2 and HOI. The plots correspond to the nighttime
averaged (from 00:00 to 01:00 LT (local time)) differences between the model
scenarios. Considerable reductions of up to 0.5 and 10 pmol mol-1
(i.e. up to 100 % removal) are observed for I2 and HOI,
respectively, particularly over coastal polluted regions where continental
pollution outflow leads to higher levels of NO3 in the nighttime MBL.
Major shipping routes also show strong nocturnal iodine activity due to the
characteristically high NOx, and resulting NO3, associated with
shipping emissions.
Figure 8 shows the effect of this nocturnal chemistry on the concentrations
of IONO2 and NO3. As in the previous figure, the plots correspond
to the nighttime averaged difference between the second and the first
scenarios. The maps show an increase of IONO2 of up to
15 pmol mol-1 (∼ 600 %) over polluted coastal areas, due to
efficient conversion of NO3 into IONO2. The bottom panel of
Figure 7 shows the expected decrease of NO3 levels associated with the
inclusion of Reactions (R1) and (R4), with decreases of up to ∼ 4 pmol mol-1 (up to 60 %) over marine polluted regions. We
model global percentage reductions in the NO3 concentrations of
7.1 % (60∘ S–60∘ N), with nitrate removal of up to
80 % in non-polluted remote oceanic regions with low NO3 levels.
This in turn can affect the modelled oxidation of DMS by NO3. We
estimate that the reduction in NO3, due to the inclusion of
Reactions (R1) and (R4), results in a model increase in DMS levels of up to
7 pmol mol-1 (about 20 %) in marine regions affected by
continental pollution outflow (Fig. 9). We therefore suggest that the
inclusion of the new nighttime iodine chemistry can have a large, so far
unrecognized, impact on the nocturnal oxidizing capacity of the marine
atmosphere.
The hourly evolution of the main species involved in this study is shown in
Figs. 10 and 11, which include the levels of HOI, I2, IONO2 and
NO3 in the MBL over regions where nocturnal iodine is modelled to be
particularly active. The first region is located within the Mediterranean
Sea, an area that shows large differences during the summer months when high
ozone levels drive large emissions of HOI and I2 from the sea, and the
high levels of NO3 at nighttime make this chemistry especially
important. The hourly average in August is shown in Fig. 10 for HOI,
IONO2 and I2. HOI and IONO2 (Fig. 10) are the species whose
concentration differ most between scenarios as HOI is removed and IONO2
produced by Reaction (R4) (and, to a lesser extent, Reaction R1). Over the Pacific Ocean at
the south of the Baja California Peninsula, the modelled differences between
the two scenarios are even higher than over the Mediterranean Sea (Fig. 11).
Large differences in MBL NO3, up to 28 %, are modelled during
the night caused by pollution outflow from the west coasts of Mexico and USA.
Summary and conclusions
The viability of the reaction of HOI with NO2, HNO3 and NO3
has been studied by theoretical calculations. The results indicate that only
the reaction of HOI with NO3, to yield IO + HNO3, is possible
under tropospheric conditions. The inclusion of this reaction, along with
that of I2+ NO3, has a number of significant implications:
Nocturnal iodine radical chemistry is activated.
This causes
enhanced nighttime oceanic emissions of HOI and I2.
Nighttime iodine species are partitioned into high levels of
IONO2.
The IO spike, modelled by current iodine models but
not shown by observations, is removed.
A reduction of the
levels of nitrate radical in the MBL, with the associated less efficient
oxidation of DMS, has important implications for our understanding of
the nocturnal oxidizing capacity of the marine atmosphere.
Data availability
Data are available upon request to the corresponding author.
The Supplement related to this article is available online at doi:10.5194/acp-16-15593-2016-supplement.
Acknowledgements
This work was supported by the Spanish National Research Council (CSIC). The
National Center for Atmospheric Research (NCAR) is funded by the National
Science Foundation (NSF). The Climate Simulation Laboratory at NCAR's
Computational and Information Systems Laboratory (CISL) provided the
computing resources (ark:/85065/d7wd3xhc). As part of the CESM project,
CAM-Chem is supported by the NSF and the Office of Science (BER) of the US
Department of Energy. This work was also sponsored by the NASA Atmospheric
Composition Modeling and Analysis Program (ACMAP, number NNX11AH90G).
Edited by: J.-U. Grooß
Reviewed by: H. K. Roscoe, F. Louis, and two anonymous referees
ReferencesAllan, B. J., McFiggans, G., Plane, J. M. C., Coe, H., and McFadyen, G. G.:
The nitrate radical in the remote marine boundary layer, J.
Geophys. Res.-Atmos., 105, 24191–24204, 10.1029/2000jd900314,
2000.Allan, J. D., Williams, P. I., Najera, J., Whitehead, J. D., Flynn, M. J.,
Taylor, J. W., Liu, D., Darbyshire, E., Carpenter, L. J., Chance, R.,
Andrews, S. J., Hackenberg, S. C., and McFiggans, G.: Iodine observed in new
particle formation events in the Arctic atmosphere during ACCACIA, Atmos.
Chem. Phys., 15, 5599–5609, 10.5194/acp-15-5599-2015, 2015.Bartis, J. T. and Widom, B.: Stochastic models of the interconversion of
three or more chemical species, J. Chem. Phys., 60, 3474–3482,
10.1063/1.1681562, 1974.Butler, J. H., King, D. B., Lobert, J. M., Montzka, S. A., Yvon-Lewis, S. A.,
Hall, B. D., Warwick, N. J., Mondeel, D. J., Aydin, M., and Elkins, J. W.:
Oceanic distributions and emissions of short-lived halocarbons, Global
Biogeochem. Cy., 21, GB1023, 10.1029/2006gb002732, 2007.
Carpenter, L. J.: Iodine In the marine Boundary Layer, Chem. Rev., 103,
4953–4962, 2003.Carpenter, L. J., Fleming, Z. L., Read, K. A., Lee, J. D., Moller, S. J.,
Hopkins, J. R., Purvis, R. M., Lewis, A. C., Müller, K., Heinold, B.,
Herrmann, H., Fomba, K. W., Pinxteren, D., Müller, C., Tegen, I.,
Wiedensohler, A., Müller, T., Niedermeier, N., Achterberg, E. P., Patey,
M. D., Kozlova, E. A., Heimann, M., Heard, D. E., Plane, J. M. C., Mahajan,
A., Oetjen, H., Ingham, T., Stone, D., Whalley, L. K., Evans, M. J., Pilling,
M. J., Leigh, R. J., Monks, P. S., Karunaharan, A., Vaughan, S., Arnold, S.
R., Tschritter, J., Pöhler, D., Frieß, U., Holla, R., Mendes, L. M.,
Lopez, H., Faria, B., Manning, A. J., and Wallace, D. W. R.: Seasonal
characteristics of tropical marine boundary layer air measured at the Cape
Verde Atmospheric Observatory, J. Atmos. Chem., 67, 87–140,
10.1007/s10874-011-9206-1, 2011.Carpenter, L. J., MacDonald, S. M., Shaw, M. D., Kumar, R., Saunders, R. W.,
Parthipan, R., Wilson, J., and Plane, J. M. C.: Atmospheric iodine levels
influenced by sea surface emissions of inorganic iodine, Nat. Geosci., 6,
108–111, 10.1038/ngeo1687, 2013.Carpenter, L. J., Andrews, S. J., Lidster, R. T., Saiz-Lopez, A.,
Fernandez-Sanchez, M., Bloss, W. J., Ouyang, B., and Jones, R. L.: A
nocturnal atmospheric loss of CH2I2 in the remote marine boundary
layer, J. Atmos. Chem., 10.1007/s10874-015-9320-6, 2015.Chambers, R. M., Heard, A. C., and Wayne, R. P.: Inorganic gas-phase
reactions of the nitrate radical: iodine + nitrate radical and iodine atom
+ nitrate radical, J. Phys. Chem., 96, 3321–3331,
10.1021/j100187a028, 1992.Fernandez, R. P., Salawitch, R. J., Kinnison, D. E., Lamarque, J.-F., and
Saiz-Lopez, A.: Bromine partitioning in the tropical tropopause layer:
implications for stratospheric injection, Atmos. Chem. Phys., 14,
13391–13410, 10.5194/acp-14-13391-2014, 2014.
Frisch, M., Trucks, G., Schlegel, H., Scuseria, G., Robb, M., Cheeseman, J.,
Scalmani, G., Barone, V., Mennucci, B., and Petersson, G.: Gaussian 09,
Revision A. 1., Wallingford, CT: Gaussian, Inc, 2009.Georgievskii, Y. and Klippenstein, S. J.: Long-range transition state theory,
J. Chem. Phys., 122, 194103, 10.1063/1.1899603, 2005.
Gilbert, R. G. and Smith, S. C.: Theory of Unimolecular and Recombination
Reactions, Blackwell, Oxford, 1990.
Glukhovtsev, M. N., Pross, A., McGrath, M. P., and Radom, L.: Extension of
Gaussian-2 (G2) theory to bromine- and iodine-containing molecules: Use of
effective core potentials, J. Chem. Phys., 103, 1878–1885, 1995.Gómez Martín, J. C., Galvez, O., Baeza-Romero, M. T., Ingham, T.,
Plane, J. M. C., and Blitz, M. A.: On the mechanism of iodine oxide particle
formation, Phys. Chem. Chem. Phys., 15, 15612–15622,
10.1039/c3cp51217g, 2013.Gómez Martín, J. C., Mahajan, A. S., Hay, T. D., Prados-Román,
C., Ordóñez, C., MacDonald, S. M., Plane, J. M. C., Sorribas, M.,
Gil, M., Paredes Mora, J. F., Agama Reyes, M. V., Oram, D. E., Leedham, E.,
and Saiz-Lopez, A.: Iodine chemistry in the eastern Pacific marine boundary
layer, J. Geophys. Res.-Atmos., 118, 887–904, 10.1002/jgrd.50132, 2013.Großmann, K., Frieß, U., Peters, E., Wittrock, F., Lampel, J.,
Yilmaz, S., Tschritter, J., Sommariva, R., von Glasow, R., Quack, B.,
Krüger, K., Pfeilsticker, K., and Platt, U.: Iodine monoxide in the
Western Pacific marine boundary layer, Atmos. Chem. Phys., 13, 3363–3378,
10.5194/acp-13-3363-2013, 2013.
Hoffmann, T., O'Dowd, C. D., and Seinfeld, J. H.: Iodine oxide homogeneous
nucleation: An explanation for coastal new particle production, Geophys. Res.
Lett., 28, 1949–1952, 2001.Holmes, N. S., Adams, J. W., and Crowley, J. N.: Uptake and reaction of HOI
and IONO2 on frozen and dry NaCl / NaBr surfaces and H2SO4,
Phys. Chem. Chem. Phys., 3, 1679–1687, 10.1039/b100247n, 2001.Jones, C. E., Hornsby, K. E., Sommariva, R., Dunk, R. M., von Glasow, R.,
McFiggans, G., and Carpenter, L. J.: Quantifying the contribution of marine
organic gases to atmospheric iodine, Geophys. Res. Lett., 37, L18804,
10.1029/2010GL043990, 2010.Kaltsoyannis, N. and Plane, J. M. C.: Quantum chemical calculations on a
selection of iodine-containing species (IO, OIO, INO3, (IO)2,
I2O3, I2O4 and I2O5) of importance in the
atmosphere., Phys. Chem. Chem. Phys., 10, 1723–1733, 2008.
Khanniche, S., Louis, F., Cantrel, L., and Černušák, I.: A
Density Functional Theory and ab Initio Investigation of the Oxidation
Reaction of CO by IO Radicals, J. Phys. Chem. A, 120, 1737–1749, 2016.Lamarque, J.-F., Emmons, L. K., Hess, P. G., Kinnison, D. E., Tilmes, S.,
Vitt, F., Heald, C. L., Holland, E. A., Lauritzen, P. H., Neu, J., Orlando,
J. J., Rasch, P. J., and Tyndall, G. K.: CAM-chem: description and evaluation
of interactive atmospheric chemistry in the Community Earth System Model,
Geosci. Model Dev., 5, 369–411, 10.5194/gmd-5-369-2012, 2012.Lawler, M. J., Mahajan, A. S., Saiz-Lopez, A., and Saltzman, E. S.:
Observations of I2 at a remote marine site, Atmos. Chem. Phys., 14,
2669–2678, 10.5194/acp-14-2669-2014, 2014.Lee, J. D., Moller, S. J., Read, K. A., Lewis, A. C., Mendes, L., and
Carpenter, L. J.: Year-round measurements of nitrogen oxides and ozone in the
tropical North Atlantic marine boundary layer, J. Geophys. Res.-Atmos., 114,
D21302, 10.1029/2009jd011878, 2009.Lee, J. D., McFiggans, G., Allan, J. D., Baker, A. R., Ball, S. M., Benton,
A. K., Carpenter, L. J., Commane, R., Finley, B. D., Evans, M., Fuentes, E.,
Furneaux, K., Goddard, A., Good, N., Hamilton, J. F., Heard, D. E., Herrmann,
H., Hollingsworth, A., Hopkins, J. R., Ingham, T., Irwin, M., Jones, C. E.,
Jones, R. L., Keene, W. C., Lawler, M. J., Lehmann, S., Lewis, A. C., Long,
M. S., Mahajan, A., Methven, J., Moller, S. J., Müller, K., Müller,
T., Niedermeier, N., O'Doherty, S., Oetjen, H., Plane, J. M. C., Pszenny, A.
A. P., Read, K. A., Saiz-Lopez, A., Saltzman, E. S., Sander, R., von Glasow,
R., Whalley, L., Wiedensohler, A., and Young, D.: Reactive Halogens in the
Marine Boundary Layer (RHaMBLe): the tropical North Atlantic experiments,
Atmos. Chem. Phys., 10, 1031–1055, 10.5194/acp-10-1031-2010, 2010.Leigh, R. J., Ball, S. M., Whitehead, J., Leblanc, C., Shillings, A. J. L.,
Mahajan, A. S., Oetjen, H., Lee, J. D., Jones, C. E., Dorsey, J. R.,
Gallagher, M., Jones, R. L., Plane, J. M. C., Potin, P., and McFiggans, G.:
Measurements and modelling of molecular iodine emissions, transport and
photodestruction in the coastal region around Roscoff, Atmos. Chem. Phys.,
10, 11823–11838, 10.5194/acp-10-11823-2010, 2010.MacDonald, S. M., Gómez Martín, J. C., Chance, R., Warriner, S.,
Saiz-Lopez, A., Carpenter, L. J., and Plane, J. M. C.: A laboratory
characterisation of inorganic iodine emissions from the sea surface:
dependence on oceanic variables and parameterisation for global modelling,
Atmos. Chem. Phys., 14, 5841–5852, 10.5194/acp-14-5841-2014, 2014.Mahajan, A. S., Oetjen, H., Saiz-Lopez, A., Lee, J. D., McFiggans, G. B., and
Plane, J. M. C.: Reactive iodine species in a semi-polluted environment,
Geophys. Res. Lett., 36, L16803, 10.1029/2009GL038018, 2009.Mahajan, A. S., Plane, J. M. C., Oetjen, H., Mendes, L., Saunders, R. W.,
Saiz-Lopez, A., Jones, C. E., Carpenter, L. J., and McFiggans, G. B.:
Measurement and modelling of tropospheric reactive halogen species over the
tropical Atlantic Ocean, Atmos. Chem. Phys., 10, 4611–4624,
10.5194/acp-10-4611-2010, 2010a.Mahajan, A. S., Shaw, M., Oetjen, H., Hornsby, K. E., Carpenter, L. J.,
Kaleschke, L., Tian-Kunze, X., Lee, J. D., Moller, S. J., Edwards, P.,
Commane, R., Ingham, T., Heard, D. E., and Plane, J. M. C.: Evidence of
reactive iodine chemistry in the Arctic boundary layer, J. Geophys.
Res.-Atmos., 115, D20303, 10.1029/2009JD013665, 2010b.Mahajan, A. S., Gómez Martín, J. C., Hay, T. D., Royer, S.-J.,
Yvon-Lewis, S., Liu, Y., Hu, L., Prados-Roman, C., Ordóñez, C.,
Plane, J. M. C., and Saiz-Lopez, A.: Latitudinal distribution of reactive
iodine in the Eastern Pacific and its link to open ocean sources, Atmos.
Chem. Phys., 12, 11609–11617, 10.5194/acp-12-11609-2012, 2012.McFiggans, G., Coe, H., Burgess, R., Allan, J., Cubison, M., Alfarra, M. R.,
Saunders, R., Saiz-Lopez, A., Plane, J. M. C., Wevill, D., Carpenter, L.,
Rickard, A. R., and Monks, P. S.: Direct evidence for coastal iodine
particles from Laminaria macroalgae – linkage to emissions of molecular
iodine, Atmos. Chem. Phys., 4, 701–713, 10.5194/acp-4-701-2004, 2004.Mečiarová, K., Šulka, M., Canneaux, S., Louis, F., and
Černušáka, I.: A theoretical study of the kinetics of the forward
and reverse reactions HI + CH3= I + CH4, Chem. Phys.
Lett., 517, 149–154, 2011.Nakano, Y., Ukeguchi, H., and Ishiwata, T.: Rate constant of the reaction of
NO3 with CH2I2 measured with use of cavity ring-down
spectroscopy, Chem. Phys. Lett., 430, 235–239,
10.1016/j.cplett.2006.09.002, 2006.
O'Dowd, C. D., Jimenez, J. L., Bahreini, R., Flagan, R. C., Seinfeld, J. H.,
Hameri, K., Pirjola, L., Kulmala, M., Jennings, S. G., and Hoffmann, T.:
Marine aerosol formation from biogenic iodine emissions, Nature, 417,
632–636, 2002.Ordóñez, C., Lamarque, J.-F., Tilmes, S., Kinnison, D. E., Atlas, E.
L., Blake, D. R., Sousa Santos, G., Brasseur, G., and Saiz-Lopez, A.: Bromine
and iodine chemistry in a global chemistry-climate model: description and
evaluation of very short-lived oceanic sources, Atmos. Chem. Phys., 12,
1423–1447, 10.5194/acp-12-1423-2012, 2012.Prados-Roman, C., Cuevas, C. A., Fernandez, R. P., Kinnison, D. E., Lamarque,
J.-F., and Saiz-Lopez, A.: A negative feedback between anthropogenic ozone
pollution and enhanced ocean emissions of iodine, Atmos. Chem. Phys., 15,
2215–2224, 10.5194/acp-15-2215-2015, 2015a.Prados-Roman, C., Cuevas, C. A., Hay, T., Fernandez, R. P., Mahajan, A. S.,
Royer, S.-J., Galí, M., Simó, R., Dachs, J., Großmann, K.,
Kinnison, D. E., Lamarque, J.-F., and Saiz-Lopez, A.: Iodine oxide in the
global marine boundary layer, Atmos. Chem. Phys., 15, 583–593,
10.5194/acp-15-583-2015, 2015b.
Read, K. A., Mahajan, A. S., Carpenter, L. J., Evans, M. J., Faria, B. V. E.,
Heard, D. E., Hopkins, J. R., Lee, J. D., Moller, S. J., Lewis, A. C.,
Mendes, L., McQuaid, J. B., Oetjen, H., Saiz-Lopez, A., Pilling, M. J., and
Plane, J. M. C.: Extensive halogen-mediated ozone destruction over the
tropical Atlantic Ocean, Nature, 453, 1232–1235, 2008.Read, K. A., Lee, J. D., Lewis, A. C., Moller, S. J., Mendes, L., and
Carpenter, L. J.: Intra-annual cycles of NMVOC in the tropical marine
boundary layer and their use for interpreting seasonal variability in CO, J.
Geophys. Res.-Atmos,, 114, D21303, 10.1029/2009jd011879, 2009.Roberston, S. H., Glowacki, D. R., Liang, C. H., Morley, C., Shannon, R.,
Blitz, M., and Pilling, M. J.: MESMER (Master Equation Solver for
Multi-Energy Well Reactions), 2008–2012: An object oriented C++ program
for carrying out ME calculations and eigenvalue-eigenvector analysis on
arbitrary multiple well systems, available at:
http://sourceforge.net/projects/mesmer (last access: 16 December 2016),
4.1 Edn., 2014.Roscoe, H. K., Jones, A. E., Brough, N., Weller, R., Saiz-Lopez, A., Mahajan,
A. S., Schoenhardt, A., Burrows, J. P., and Fleming, Z. L.: Particles and
iodine compounds in coastal Antarctica, J. Geophys. Res.-Atmos., 120,
7144–7156, 10.1002/2015jd023301, 2015.Saiz-Lopez, A. and Plane, J. M. C.: Novel iodine chemistry in the marine
boundary layer, Geophys. Res. Lett., 31, L04112, 10.1029/2003GL019215,
2004.Saiz-Lopez, A., Plane, J. M. C., Mahajan, A. S., Anderson, P. S., Bauguitte,
S. J.-B., Jones, A. E., Roscoe, H. K., Salmon, R. A., Bloss, W. J., Lee, J.
D., and Heard, D. E.: On the vertical distribution of boundary layer halogens
over coastal Antarctica: implications for O3, HOx,
NOx and the Hg lifetime, Atmos. Chem. Phys., 8, 887–900,
10.5194/acp-8-887-2008, 2008.Saiz-Lopez, A., Lamarque, J.-F., Kinnison, D. E., Tilmes, S.,
Ordóñez, C., Orlando, J. J., Conley, A. J., Plane, J. M. C., Mahajan,
A. S., Sousa Santos, G., Atlas, E. L., Blake, D. R., Sander, S. P.,
Schauffler, S., Thompson, A. M., and Brasseur, G.: Estimating the climate
significance of halogen-driven ozone loss in the tropical marine troposphere,
Atmos. Chem. Phys., 12, 3939–3949, 10.5194/acp-12-3939-2012, 2012a.Saiz-Lopez, A., Plane, J. M. C., Baker, A. R., Carpenter, L. J., Von Glasow,
R., Gómez Martín, J. C., McFiggans, G., and Saunders, R. W.:
Atmospheric Chemistry of Iodine, Chem. Rev., 112, 1773–1804,
10.1021/cr200029u, 2012b.Saiz-Lopez, A., Fernandez, R. P., Ordóñez, C., Kinnison, D. E.,
Gómez Martín, J. C., Lamarque, J.-F., and Tilmes, S.: Iodine
chemistry in the troposphere and its effect on ozone, Atmos. Chem. Phys., 14,
13119–13143, 10.5194/acp-14-13119-2014, 2014.Saiz-Lopez, A., Baidar, S., Cuevas, C. A., Koenig, T. K., Fernandez, R. P.,
Dix, B., Kinnison, D. E., Lamarque, J. F., Rodriguez-Lloveras, X., Campos, T.
L., and Volkamer, R.: Injection of iodine to the stratosphere, Geophys. Res.
Lett., 42, 6852–6859, 10.1002/2015gl064796, 2015.
Sander, S. P., Friedl, R. R., Golden, D. M., Kurylo, M. J., Moortgat, G. K.,
Wine, P. H., Ravishankara, A. R., Kolb, C. E., Molina, M. J., Diego, S.,
Jolla, L., Huie, R. E., and Orkin, V. L.: Chemical Kinetics and Photochemical
Data for Use in Atmospheric Studies Evaluation Number 15, JPL_NASA, 06-2,
Jet Propulsion Laboratory, Pasadena, CA, 2006.Sherwen, T., Evans, M. J., Carpenter, L. J., Andrews, S. J., Lidster, R. T.,
Dix, B., Koenig, T. K., Sinreich, R., Ortega, I., Volkamer, R., Saiz-Lopez,
A., Prados-Roman, C., Mahajan, A. S., and Ordóñez, C.: Iodine's
impact on tropospheric oxidants: a global model study in GEOS-Chem, Atmos.
Chem. Phys., 16, 1161–1186, 10.5194/acp-16-1161-2016, 2016.Simpson, W. R., Brown, S. S., Saiz-Lopez, A., Thornton, J. A., and Glasow, R.
v.: Tropospheric Halogen Chemistry: Sources, Cycling, and Impacts, Chem.
Rev., 115, 4035–4062, 10.1021/cr5006638, 2015.Sommariva, R., Bloss, W. J., and von Glasow, R.: Uncertainties in gas-phase
atmospheric iodine chemistry, Atmos. Environ., 57, 219–232,
10.1016/j.atmosenv.2012.04.032, 2012.
Šulková, K., Šulka, M., Louis, F., and Neogrády, P.:
Atmospheric Reactivity of CH2ICl with OH Radicals: High-Level OVOS
CCSD(T) Calculations for the XAbstraction Pathways (X = H, Cl, or I), J.
Phys. Chem. A, 117, 771–782, 2013.Volkamer, R., Baidar, S., Campos, T. L., Coburn, S., DiGangi, J. P., Dix, B.,
Eloranta, E. W., Koenig, T. K., Morley, B., Ortega, I., Pierce, B. R.,
Reeves, M., Sinreich, R., Wang, S., Zondlo, M. A., and Romashkin, P. A.:
Aircraft measurements of BrO, IO, glyoxal, NO2, H2O, O2–O2 and
aerosol extinction profiles in the tropics: comparison with
aircraft-/ship-based in situ and lidar measurements, Atmos. Meas. Tech., 8,
2121–2148, 10.5194/amt-8-2121-2015, 2015.von Glasow, R., Sander, R., Bott, A., and Crutzen, P. J.: Modeling halogen
chemistry in the marine boundary layer. 1. Cloud-free MBL, J. Geophys. Res.,
107, 4341, 10.1029/2001JD000942, 2002.Wachsmuth, M., Gäggeler, H. W., von Glasow, R., and Ammann, M.:
Accommodation coefficient of HOBr on deliquescent sodium bromide aerosol
particles, Atmos. Chem. Phys., 2, 121–131, 10.5194/acp-2-121-2002, 2002.Wang, F., Saiz-Lopez, A., Mahajan, A. S., Gómez Martín, J. C.,
Armstrong, D., Lemes, M., Hay, T., and Prados-Roman, C.: Enhanced production
of oxidised mercury over the tropical Pacific Ocean: a key missing oxidation
pathway, Atmos. Chem. Phys., 14, 1323–1335, 10.5194/acp-14-1323-2014,
2014.