ACPAtmospheric Chemistry and PhysicsACPAtmos. Chem. Phys.1680-7324Copernicus GmbHGöttingen, Germany10.5194/acp-15-9731-2015A mechanism for biologically induced iodine emissions from sea iceSaiz-LopezA.a.saiz@csic.eshttps://orcid.org/0000-0002-0060-1581Blaszczak-BoxeC. S.CarpenterL. J.Atmospheric Chemistry and Climate Group, Institute of Physical
Chemistry Rocasolano, CSIC, Madrid, SpainDepartment of Physical, Environmental and Computer Sciences, Medgar
Evers College-City University of New York, Brooklyn, NY 11235, USACUNY Graduate Center, Chemistry Division, Earth and Environmental
Science Division, Manhattan, NY 10016, USAWolfson Atmospheric Chemistry Laboratories, Department of Chemistry,
University of York, Heslington, York YO10 5DD, UKA. Saiz-Lopez (a.saiz@csic.es)1September201515179731974612February20158April201514July201520July2015This 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/15/9731/2015/acp-15-9731-2015.htmlThe full text article is available as a PDF file from https://acp.copernicus.org/articles/15/9731/2015/acp-15-9731-2015.pdf
Ground- and satellite-based measurements have reported high concentrations of
iodine monoxide (IO) in coastal Antarctica. The sources of such a large
iodine burden in the coastal Antarctic atmosphere remain unknown. We propose
a mechanism for iodine release from sea ice based on the premise that
micro-algae are the primary source of iodine emissions in this environment.
The emissions are triggered by the biological production of iodide (I-)
and hypoiodous acid (HOI) from micro-algae (contained within and underneath
sea ice) and their diffusion through sea-ice brine channels, ultimately accumulating in a thin
brine layer (BL) on the surface of sea ice. Prior to reaching the BL, the
diffusion timescale of iodine within sea ice is depth-dependent. The BL is
also a vital component of the proposed mechanism as it enhances the chemical
kinetics of iodine-related reactions, which allows for the efficient release
of iodine to the polar boundary layer. We suggest that iodine is released to
the atmosphere via three possible pathways: (1) emitted from the BL and then
transported throughout snow atop sea ice, from where it is released to the
atmosphere; (2) released directly from the BL to the atmosphere in regions of
sea ice that are not covered with snowpack; or (3) emitted to the atmosphere
directly through fractures in the sea-ice pack. To investigate the proposed
biology–ice–atmosphere coupling at coastal Antarctica we use a multiphase
model that incorporates the transport of iodine species, via diffusion, at
variable depths, within brine channels of sea ice. Model simulations were
conducted to interpret observations of elevated springtime IO in the coastal
Antarctic, around the Weddell Sea. While a lack of experimental and
observational data adds uncertainty to the model predictions, the results
nevertheless show that the levels of inorganic iodine (i.e. I2, IBr,
ICl) released from sea ice through this mechanism could account for the
observed IO concentrations during this timeframe. The model results also
indicate that iodine may trigger the catalytic release of bromine from sea
ice through phase equilibration of IBr. Considering the extent of sea ice
around the Antarctic continent, we suggest that the resulting high levels of
iodine may have widespread impacts on catalytic ozone destruction and aerosol
formation in the Antarctic lower troposphere.
Introduction
Over the past two decades, evidence has accumulated for the role of
atmospheric iodine in the catalytic destruction of tropospheric ozone (e.g.
Chameides and Davis, 1980; Solomon et al., 1994; Vogt et al., 1999; McFiggans
et al., 2000; Calvert and Lindberg, 2004a; Saiz-Lopez et al., 2007a, 2012a,
2014; Read et al., 2008; Sommariva and von Glasow, 2012; Carpenter et al.,
2013). In the mid-latitude marine boundary layer (MBL) iodine decreases the
HOx ratio (i.e. [HO2] / [OH]) via the reaction of iodine
monoxide (IO) and HO2 to form hypoiodous acid (HOI), which then
photolyses efficiently to OH, whereas NO oxidation to NO2 by IO
increases the NOx ratio (i.e. [NO2] / [NO]) (e.g. Bloss et al.,
2005, 2010). Iodine atoms released from photolysis react
rapidly with O3 to form IO; rapid reactions of IO with HO2 and
NO2, followed by photochemical breakdown, result in the regeneration of
atomic iodine without forming O3 and therefore in catalytic O3
destruction.
Considerable attention has been given to the role of iodine oxides in the
formation of ultra-fine aerosol (i.e. 3–10 nm diameter) and its potential
to contribute to cloud condensation nuclei (CCN) formation (O'Dowd et al.,
1998, 2002; Hoffmann et al., 2001; Jimenez et al., 2003; Saiz-Lopez and
Plane, 2004; McFiggans et al., 2004; Burkholder et al., 2004; Sellegri et
al., 2005; Saunders and Plane, 2005, 2006; Saiz-Lopez et al., 2006; Pechtl et
al., 2006; Atkinson et al., 2012; Saiz-Lopez et al., 2012b; Gómez
Martín et al., 2013; Gálvez et al., 2013). Iodine has also been
suggested to impact the depletion of gaseous elemental mercury (Hg∘)
by oxidation to reactive gaseous mercury (HgII) (Calvert and Lindberg,
2004b; Saiz-Lopez et al., 2008; Wang et al., 2014).
IO is formed following photolysis of photo-labile reactive iodine precursors
and the subsequent reaction of I atoms with atmospheric O3. In the polar
boundary layer, IO has been detected in coastal Antarctica by ground-
(Friess et al., 2001; Saiz-Lopez et al., 2007a, Atkinson et al., 2012) and
satellite-based instrumentation (Saiz-Lopez et al., 2007b; Schönhardt et
al., 2008, 2012), and also by ground-based techniques in the Arctic boundary
layer (Mahajan et al., 2010). These studies have shown iodine to be very
abundant (e.g. up to 20 pptv during Antarctic springtime) and widespread
around coastal Antarctica, which have been proposed to significantly impact
the chemistry and vertical distribution of O3, HOx, NOx, and
Hg in the coastal Antarctic marine boundary layer (Saiz-Lopez et al., 2008).
In the polar regions the source of reactive inorganic bromine and chlorine
includes heterogeneous reactions involving sea-salt bromide on sea ice,
snowpack, or marine aerosol surfaces (e.g. Saiz-Lopez and von Glasow, 2012,
and references therein). These heterogeneous reactions take part in an
autocatalytic cycle that destroys ozone while preserving atomic halogen
radicals. However, these mechanisms do not result in significant iodine
release to the gas phase due to the comparatively much smaller content of
iodide (I-) in sea salt. Over the continental Antarctic snowpack two
mechanisms have been proposed to explain the observed large levels of IO: (i)
direct release of reactive iodine from snowpack (Frieß et al., 2010) and
(ii) aerosol deposition and subsequent recycling on snowpack (Saiz-Lopez et
al., 2008). However, in sea-ice areas around coastal Antarctica, iodocarbon
(i.e. CH3I, CH2ICl, CH2I2, CH2IBr) levels have been
found to be insufficient to account for the high concentrations of gas-phase
iodine due to their relatively long photolytic lifetimes, as compared to
inorganic precursors, and small concentrations (Atkinson et al., 2012;
Granfors et al., 2015). Therefore, although the presence of high levels of
reactive iodine in the coastal Antarctic boundary layer has already been
evidenced, the sources and mechanisms of iodine release over sea-ice-covered
areas in coastal Antarctica still remain unknown. In this study a hypothesis
for iodine release from sea ice is presented. The suggested coupling between
biology, sea ice and overlying atmosphere is investigated using a multiphase
chemical model.
Physical context of the polar sea-ice environment
Sea ice is one of the most extreme and largest ecosystems in the polar ocean
(Eicken, 1992a, 2003a; Brierley and Thomas, 2002; Arrigo and Thomas, 2004).
Sea ice is not completely solid and can, at times, be comprised of a system
of brine channels that provide a habitat, characterized by low temperatures
(253 ≤ T/K ≤ 271), high salinity (35–200 psu), high pH (up to
11), and low irradiances (1 µmol photons m-2 s-1) at the
sea-ice–water interface (Eicken, 1992b; Gleitz et al., 1995; Kirst and
Wiencke, 1995).
Seawater starts to freeze as temperatures drop below -1.86 ∘C
(271.30 K) since it generally contains about 35 g L-1 (salinity of 3.5 %) of dissolved salts (e.g. sodium, calcium, sulfate, magnesium,
chloride, and potassium). Ice begins to form and rise to the surface, forming
frazil ice in various physical shapes and dimensions. Thereafter, loosely
aggregated pancakes (ice discs) form by the motion of wind and water to
consolidate ice crystals. These pancake ice features propagate in both the
lateral and vertical direction as time progresses, eventually forming packed
ice, which is permeable to microscopic transport and impermeable to
macroscopic transport – especially macroscopic transport of brine channel
fluid. The vast majority of Antarctic sea ice is comprised of frazil and
platelet ice, compared to 80 % of columnar ice in the Arctic; this is
primarily due to the fact that Antarctic sea ice is formed with more
turbulent water compared to much calmer conditions experienced in the Arctic
(Margesin et al., 2007; Eicken, 2003b).
Salts, ions, and air in the water are concentrated into inclusions of pockets
and channels or released into the water below the sea ice as they cannot be
incorporated into ice crystals during the formation of sea ice. Hence, sea
ice is a solid matrix penetrated by a labyrinth of channels and pores that
contain highly concentrated brine and air bubbles. Brine channels are defined
as vertically elongated tubular systems containing fluid, with diameters of
less than a few millimetres to several centimetres and exhibit a vertical
extent up to 50 cm at coastal Antarctica (Thomas and Dieckmann, 2003). They
are also the main habitat for all micro-organisms in sea ice (Brierley and
Thomas, 2002; Deming, 2002; Lizotte, 2003; Mock and Thomas, 2005; Mock and
Junge, 2007). Salt concentration and brine volume are directly proportional
to temperature (Eicken, 2003; Weissenberger et al., 1992; Krembs et al.,
2000); therefore, as temperature increases, brine volume increases and salt
concentration decreases. Hence, colder ice contains brine channels with
highly salty brines and overall fewer, smaller and less interconnected
channels than warmer ice. Gradients exist in temperature, brine salinity, and
brine volume since ice at the air–ice interface is usually colder than ice
in contact with underlying water. A suite of protists and zooplankton have
been recorded to live in sea ice (Horner, 1985; Palmisano and Garrison, 1993;
Lizotte, 2003; Schnack-Schiel, 2003; Werner, 2006). Among the autotrophs, the
most studied are diatoms. All discovered organisms within sea ice have
plastic physiologies to cope with these dynamic changes (dominated by
temperature and salinity changes) in physical and chemical conditions of
their environment.
As sea ice forms, micro-algae get caught between the ice crystals or simply
stick to them as crystals rise through the water when it freezes in the
autumn. Diatoms become trapped within brine channels during the formation of
consolidated ice. Pennate diatoms, along with other micro-algae (e.g.
dinoflagellates, flagellates), are the most conspicuous organisms in sea
ice (Brierley and Thomas, 2002; Thomas and Dieckmann, 2002; Lizotte, 2003).
These micrometre-sized algae, with their main light-harvesting pigment,
fucoxanthin, can attain such concentrations in sea ice that they discolour
the ice visibly brown. The time for acclimation to new conditions is not very
long since daylight hours continue to decrease as winter approaches.
Nevertheless, diatoms, especially at the ice–water interface, are often able
to photo-acclimate rapidly and can accumulate to high biomass even before the
winter begins (Gleitz and Thomas, 1993). Sea-ice diatoms are very efficient
at optimizing solar irradiance and are able to grow at low irradiance levels
– below 1 µmol photons m-2 s-1 (Mock and Graddinger,
1999). Light levels are minimal during polar winter due to short
days/complete darkness and snow cover atop sea ice, which acts as a very
efficient reflector of solar irradiance (Eicken, 2003).
Maximum growth rates for polar diatoms are 0.25–0.75 divisions per day –
i.e. 2- to 3-fold slower than growth rates at temperatures above
10 ∘C (Sommer, 1989). Many of these diatoms are psychrophilic and
cannot live at temperatures above ca. 15 ∘C, indicative of the
presence of specific molecular adaptations that enable these diatoms to grow
under freezing temperatures. Only recently, functional genomics were applied
to various/specific diatoms to begin to uncover the molecular basis of growth
and, thus, adaptation to polar conditions (Mock and Valtentin, 2004; Mock et
al., 2006).
Most notably, Hill and Manley (2009) investigated the release of reactive
iodine from diatoms. Via an in situ incubation assay, they measured the
iodination of phenol red to detect the release of reactive iodine (primarily
HOI) from a putative extracellular bromoperoxidase of marine
diatoms. Six of 11 species showed significant release compared to controls.
Polar diatoms were especially active, releasing 0.02–2.7 µmol HOI
[mg total chlorophyll]-1 h-1, at 100 µmol L-1
iodide concentration. Therefore, micro-algae are a major source of iodine,
complementing the study of Küpper et al. (1998) of enhanced iodine uptake in
macro-algae. Hill and Manley (2009) find that the release of not only
iodine but also bromine species can account for a significant amount of
their emissions needed to simulate polar tropospheric ozone depletion events.
Model description
In order to study the link between polar marine micro-algae iodine emissions
and the potential for iodine release from sea ice, we developed the
multiphase chemical model (Condensed Phase to Air Transfer Model,
CON-AIR). This model incorporates the multi-component aspect of sea ice
(e.g. ice, quasi-liquid layer (brine layers – BLs), brine channels, and micro-algae),
interfaced with overlying atmospheric boundary layer chemistry. It is
structured in three main components: (i) micro-algae emissions and transport
through sea-ice brine channels; (ii) aqueous-phase chemistry regime in the
BL; and (iii) gas-phase chemistry scheme comprising photochemical, thermal
and heterogeneous reactions.
Aqueous- and gas-phase chemistry in CON-AIR
The BL is defined as a thin layer on the surface of sea ice and ice crystals, where
water molecules are not in a rigid, solid structure, yet not in the random
order of liquid (Petrenko and Whitworth, 1999). The aqueous-phase component
treats 14 species and comprises 25 condensed-phase reactions, representing
iodine, bromine, and chlorine chemistry in the BL. The gas-phase chemistry
includes 41 chemical species, 154 reactions representative of the standard
O3–NOx–HOx–S and halogen gas-phase chemistry, along with a
treatment of the halogen recycling on deliquesced airborne sea-salt aerosols.
It also incorporates 12 processes of heterogeneous uptake and wet/dry
deposition and 38 photochemical reactions. The complete scheme of reactions
in the BL and the gas phase employed in the model is summarized in the
Supplement.
The exchange of halogen species between the liquid (BL) and gas phase is
treated via phase equilibration as well as deposition of gas-phase molecules
onto the sea-ice surface. For iodine species, this liquid–gas-phase exchange
depends upon the Henry's law constants of species including I2, IBr and
ICl. The gas-phase equilibrating species, Henry's law solubility constants,
and temperature dependences are shown in Table 3 of the Supplement. The temperature dependence of the solubility of species is taken
into account by including a diurnal variation of the typical temperature
profile during springtime (i.e. ∼ 260 ≤ T/K ≤∼ 270) (Launiainen and Vihma, 1994).
Transport through sea ice
Here, we consider only transport by diffusion, but this should be regarded as
a lower limit as transport is also governed by wind pumping, advection,
thermal convection, and/or fluid transport. Fick's first law is used in
steady-state diffusion – i.e. when the concentration within the diffusion
volume does not change with respect to time (Jin=Jout). In one (spatial) dimension,
J=-D∂ϕ∂x,
where J is the diffusion flux (amount of
substance/length2 time-1), D is the diffusion coefficient or
diffusivity (length2/time), ϕ is the concentration (amount of
substance/volume), and x is the position (length). The range of diffusion
coefficients used in this study are from 10-4 to
10-7 cm2 s-1, which are within the range of experimental
data on diffusion of species in ice/snow (Shaw et al., 2011; Loose et al.,
2011; Callaghan et al., 1999; Mercier et al., 2005). We also use Fick's
second law as the transport of iodine is in non-steady state (or continually
changing) since the concentration within the diffusion volume changes with
respect to time.
∂ϕ∂t=D∂2ϕ∂x2,
where t is time.
Here, we solve Fick's second law via the limited-source-diffusion
approximation and incorporate this solution into the model. Given the
following boundary conditions,
φ(x,0)=0∫φ(x,t)dx=Sφ(x,∞)=0,
where S is called the “dose”, where S(t)=φo(4Dt/π)1/2, where φo
is the initial concentration of iodine at the surface in the brine layer. The
solution to Fick's law under these conditions is φ(x,t)= (S/(πDt)1/2) × exp(-x2/4Dt).
Therefore, at each time “t” ``φ(x,t)” was computed and then
incorporated in Fick's first law at each specified time “t” to compute J at each specified time “t”. We used this approximation to take into
account the change in J as a function of time.
From late July/early August (transition from winter to spring) till late
November/early December (transition from spring to summer), coastal Antarctic
temperatures range from ∼ 208 to 273 K (Schwerdtfeger, 1974;
Veihelmann et al., 2001). This temperature regime encompasses both
microscopic (gaseous diffusion through the snowpack/sea-ice crystal network,
transport through water/brine veins) and macroscopic (transport of bulk brine
through brine channels, water-vein transport) transport through sea ice (see
Sect. 2). This timeframe also coincides with the onset of the release of
iodine and gradual decline of iodine release, from the start of summer
onwards (Saiz-Lopez et al., 2007a). The Antarctic Peninsula and the
temperature regime of the Weddell Sea exhibit a temperature range from 256 to
270 K from the beginning until the end of spring (September to December).
Via monthly averaged surface temperatures (Schwerdtfeger, 1974), the
Antarctic Peninsula's east coast experiences temperatures 8 ∘C
colder than the Antarctic Peninsula's west coast. During the timeframe of
iodine release, the Antarctic Peninsula's west coast experiences temperatures
at or above -5 ∘C, the lower-limit temperature demarcation,
governing the “rule of fives” and thus lower-limit temperature boundary
for macroscopic permeability (Golden et al., 1998). Prior to this study, it
was shown that (unlike freshwater) sea ice is a highly permeable medium for
gases. It was shown that the migration of gases along grain boundaries (e.g.
in brine channels) was 2 to 6 times as great as that at right angles to the
principal axis of the grain boundaries (Gosink et al., 1976). At
-15 ∘C, penetration rates of halogenated species were 30 and
60 cm h-1 for CO2 (Gosink et al., 1976). The vertical migration
is approximately twice as fast as horizontal migration at 15 ∘C; for
semi-fresh pressure ridge ice, the migration rate was 60 cm h-1 at
-15 ∘C. Gosink et al. (1976) produced estimated permeation
constants of 10-7 and 10-5 cm2 s-1 atm-1 for
SF6 at -15 ∘C and CO2 at -7 ∘C.
Golden et al. (2007) showed that the columnar sea-ice permeability for
liquids drop by approximately 2 orders of magnitude below a 5 %
relative brine volume, which corresponds to roughly -5 % for a bulk
salinity of 5 (i.e. the “rule of fives”). It has long been observed
(Weeks and Ackley, 1986) that columnar sea ice is effectively impermeable to
brine transport for ice porosity less than 5 %, yet is permeable for ice
porosity above 5 %. For a bulk salinity of 5 parts per thousand (ppt),
the critical ice porosity ∼ 5 % corresponds to a critical
temperature of -5 ∘C, via equations relating ice porosity to
temperature and salinity (Thomas and Diekmann, 2003; Weeks and Ackley, 1986).
Golden et al. (1998) discussed the rule of fives in terms where the critical
ice porosity was identified with the critical probability in a continuum
percolation model for a compressed powder (Kusy and Turner, 1971) –
exhibiting microstructural characteristics qualitatively similar to sea ice.
In the Arctic, strongly aligned columnar ice is the dominant textural type
and accounts for approximately two-thirds to three-quarters of the total ice volume. Dynamic growth
conditions in the Antarctic limit the occurrence of columnar ice to the
lowermost layers of the ice cover. While vertically oriented columnar
crystals are common, horizontal alignment is observed only infrequently and
generally both horizontal and vertical dimensions of columnar crystals in
Antarctic sea ice are smaller than their Arctic counterparts (Thomas and
Diekmann, 2003). Therefore, the rule of fives has partial (at most 25 %)
and minimal application for Arctic and Antarctic sea ice, respectively, when
solely considering ice microstructure. Antarctica's ice volume is comprised
of frazil and platelet ice (Margesin et al., 2007; Eicken, 2003; Thomas and
Diekmann, 2003).
Brine entrapped in sea ice will always be at or near freezing since any
departure will either cause some of the water in the brine to freeze or some
of the surrounding ice to melt. Thus, brine salinity is variable and can be
determined based strictly on temperature. Tucker et al. (1993), Heygster et al. (2009) and Ulaby
et al. (1986) derived empirical formulas relating sea-ice temperature and
brine salinity. These equations show that from -5 to -20 ∘C
brine salinity ranges from ∼ 85 to 210 ppt. Note that the eutectic
temperature for NaCl is -21.2 ∘C; therefore, thin liquid films
should exist well below zero with the porous component of sea ice (apart from
brine channels). Additional solutes will lower the freezing point of
interfacial thin films. However, (when compared to Arctic sea ice) brine
salinity is greater (overall) in Antarctica sea ice; still, Arctic sea-ice
brine salinity is predominantly over 5 ppt (Vancoppenolle et al., 2009;
Gleitz et al., 1995; Eicken, 2012; Arrigo and Sullivan, 1992). In addition
(apart from past studies on sea-ice permeability/transport) (Gosink et al.,
1976; Boxe, 2005), recent field observations of fluxes of trace gases have
also questioned the impermeability of sea ice during winter (Heinesch et al.,
2009; Miller et al., 2011) and spring (Semiletov et al., 2004; Delille, 2006;
Zemmelink et al., 2006; Nomura et al., 2010a, b; Papakyriakou and Miller,
2011; Grannas et al., 2007).
Within the degree of uncertainty of a limited number of experiments on
species transport in snow/ice, overall it appears that species in ice have
some degree of appreciable mobility, which causes them to be impactful in
polar tropospheric boundary layer chemistry (Boxe et al., 2003, 2005, 2006;
Boxe and Saiz-Lopez, 2008; Grannas et al., 2007; Granfors et al., 2015). In
other words, especially within the context of the suite of impurities
contained in sea ice/snowpack (with varying range of concentrations),
macroscopic and microscopic transport is still possible outside of the
physical parameters that govern the “rule of fives” as exemplified
by the suite of field studies that have measured trace gases over the Arctic
and Antarctic snowpack and sea ice – over many decades – both in situ and
remotely at much lower temperatures and a wide range of salinity levels. A
given brine volume fraction can be attained by a variety of temperature and
salinity combinations as shown by the Frankestein and Garner equations
(Frankestein and Garner, 1967). Pinpointing the critical conditions for
impermeability is crucial. So, if the permeability of sea ice is so
important, why have there not been extensive permeability measurements,
like those done for porous materials? In most materials where permeability
measurements are made, the matrix material does not react with the fluid
passing through it. This is definitely not the case with sea ice, where
slight differences between the temperature of the ice matrix and of
temperature and salinity of brine can result in either the addition or the
subtraction of ice from the matrix during the experimental procedure (Ono and
Kasai, 1985; Saito and Ono, 1978; Maksym and Jeffries, 2000).
Aqueous-phase scheme and BL parameterizations
The initial concentrations of I-, Br- and Cl- in the BL are
assumed to be that of the ions in seawater – 1.3 × 10-7,
8 × 10-4 and 0.545 M, respectively (Wayne, 2000). This model
assumes that all ions and molecular species reside in the BL. In order to
account for the concentration effect on the aqueous-phase reaction rates, the
volume of the BL needs to be calculated. Using a mean thickness for the
Southern Ocean sea ice and density of 50 cm and 0.91 g cm-3 (Thomas
and Dieckmann, 2003), respectively, the total potential liquid content in a
snow column of 1 cm2 cross-sectional area of sea ice is
totalpotentialliquidcontent=50cm× 0.91gcm-31g cm-3= 45.5cm3cm-2.
The mean mass fraction of liquid water in ice between 265 and 250 K is
1 × 10-3 (Conklin and Bales, 1993). We calculate a mean BL
thickness of 500 µm using the following equation: sea-ice
thickness × sea-ice cross-sectional area × mass fraction of
liquid water = 50 cm × 1 cm2× 10-3=0.05 cm3, and then 0.05 cm3/1 cm2= 500 µm.
Still, we do acknowledge that the range of thickness of is highly variable as
shown by several experimental and modelling studies (Huthwelker et al.,
2006). The BL volume in sea ice can now be calculated as
BLvolume= 45.5cm3cm-2× 1× 10-3= 0.0455cm3cm-2.
For an atmospheric boundary layer height of 400 m (40 000 cm) a volumetric
factor is obtained:
volumetric=0.0455cm3cm-240 000cm= 1.14× 10-6(cm3(BL)cm3(atmosphere)).
Therefore, the reaction rates are quantified incorporating the volumetric
factor. We find that this enhancement in model concentrations and reaction
rates due to the concentration effect of ions and molecular species in the BL
is necessary to provide ample gas-phase concentrations. The rate constants
for the BL reactions are then expressed as
k[1cm3(atmosphere)]/[1.14× 10-6(BL)],k[1cm3(atmosphere)]2/[1.14× 10-6(BL)]2,
where k are the literature aqueous-phase rate constants in units
of cm3molecule-1s-1
and
cm6molecule-2s-1,
for second- and third-order rate constants, respectively.
The rate of transfer of species from the BL to the gas phase is calculated
using an approximation for the first-order rate constant,
kt = 1.25× 10-5 s-1, previously suggested
by Gong et al. (1997) and Michalowski et al. (2000):
kmix=kt×40 000cm3(atmosphere)0.0455cm3(BL).
However, the rate of transfer of species will depend on the concentration and
Henry's law constants for solubility of the corresponding species. Hence, the
complete expression for the phase equilibration of species from the BL to the
atmosphere is
k(BL→Atmosphere)=(kmix×[species concentration]×volumetric)/(H′),
where H′ is the dimensionless Henry's law constant. H′ is accordingly
defined as H′= (HRT), where H is a species' Henry's law constant, R is
the gas constant, 0.082058 L atm K-1 mol-1, and T is the
temperature (K). The dependence of the Henry's law constants on the
salinity was not considered due to the lack of the experimental data.
A simplified scheme of iodine cycling in and over
Antarctic sea ice. The
biological release of iodine into brine channels occurs in and on the underside of Antarctic sea ice.
Subsequently,
diffusion of iodine through brine channels allows for the accumulation of
these species in the BL, releasing I2(g) to the atmosphere via
gas-phase equilibration. Thereafter transformations of compounds occur in the
gas phase, as well as in deliquesced sea-salt aerosol.
Radiation and gas-phase scheme
Photolysis rates are calculated offline from reported absorption
cross sections and quantum yields using a two-stream radiative transfer code
(Thompson, 1984), where the irradiance reaching the surface is computed after
photon attenuation, by means of aerosol scattering and molecular absorption, through
fifty 1 km layers in the atmosphere. The model is run with surface albedo of
0.85, typical of measurements made with an actinic flux spectrometer at
Halley Station (Jones et al., 2008). The aerosol profile used in the
radiative transfer code is consistent with aerosol loadings and surface area
typical of remote locations (i.e. 10-7 cm2 cm-3).
Some species in the model are constrained to their typical values measured
during the CHABLIS (Chemistry of the Antarctic Boundary Layer and Interface with Snow) measurement field campaign that took place at Halley Bay in
coastal Antarctica (Jones et al., 2008; Read et al., 2007), with diurnal
mixing ratio profiles peaking at [CO] = 35 ppb, [DMS] = 80 pptv,
[SO2] = 15 pptv, [CH4] = 1750 ppb, [CH3CHO] = 150 pptv,
[HCHO] = 150 pptv, [isoprene] = 60 pptv, [propane] = 25 pptv,
and [propene] = 15 pptv. During the model simulations all other species are
allowed to vary. The model is solved using a variable-step-size fourth-order
Runge–Kutta integrator.
The heterogeneous recycle rate of a species on airborne sea-salt aerosols is
calculated using the free molecular transfer approximation kt= 0.25 γcA, where γ are the uptake coefficients whose
values for the different species are taken from Atkinson et al. (2006), c
is the root-mean-square molecular speed, and A is the effective available
surface area, 10-7 cm2 cm-3 (von Glasow et al., 2002) chosen
to be typical of remote oceanic conditions. The dry deposition of a species
i is computed as ViCi(t)/H, where C is the concentration
of a gaseous species at a given time and Vi is the deposition velocity
of species over a fixed boundary layer over time with a depth H of 400 m.
Typical deposition velocities in the model are 0.5 cm s-1, and we
assume the surface is flat. There is now a strong dependence of IBr release
upon the deposition velocities of HOI, HI and IONO2 since most of the
iodine present in the sea-ice brine layer comes from below via diffusion
through the brine channels.
A simplified quantitative schematic of the proposed
mechanism and the CON-AIR model structure. Note that the dimensions are not
at real scale.
Proposed mechanism for iodine release from sea ice
The mechanism is broadly illustrated in Fig. 1. Briefly, the process
includes release of iodine, in the equilibrium form of
HOI+I-+H+↔I2+H2O, from sea-ice algae and, thereafter, diffusion
through brine channels to accumulate in the BL of the ice surface accompanied
by deposition and recycling of atmospheric iodine species on the BL. Note
that, although we focus on inorganic iodine, organic iodine would also be
transported through brine channels. Figure 2 shows a simplified quantitative
schematic of the mechanism and model structure. The mechanism is based on
three characteristics that have been separately reported to occur in the
Antarctic springtime sea-ice environment:
The Southern Ocean contains the largest quantity of
micro-algae/diatoms (phytoplankton bloom) in the world (Thomas and Dieckmann,
2003). Antarctic sea ice covers an extensive portion of the Earth's surface
– that is, a maximum extent of
∼ 4 % = 19 × 106 km2 in winter and minimum
extent of ∼ 1 % = 5 × 106 km2 in summer and
is accompanied to a significant degree by biological activity; it therefore
represents one of the principal biomes on Earth (Thomas and Dieckmann, 2003).
It is known that micro-algae populations colonize the underside of sea ice
(at the seawater–sea-ice interface) and within the brine channels up to the
top of sea-ice column (Thomas and Dieckmann, 2003). Via pre-concentration
processes, these organisms contain enhanced concentrations of iodine up to
1000 times (micro-algae, e.g. Porosira glacialis/Achnanthes cf.
longipes; Hill and Manley, 2009) and 30 000 times (macro-algae, e.g.
Laminaria digitata; Küpper et al., 1998) the iodine levels
in the surrounding seawater (e.g.
[I-]seawater∼ 10-7 M). Iodide (I-)
accumulates to these high concentrations by way of a facilitated-diffusion
process, by which efficient transport and iodine uptake from natural seawater
into macro- and microalgal cells occurs, independent of its electrochemical
potential gradient (Küpper et al., 1998). Additionally, in the
extracellular domain, haloperoxidases, membrane-bound enzymes or cell wall
oxidases, along with probable intracellular sources, produce a constant flow
of H2O2 in the apoplast of cells. Therefore, within
the cells, haloperoxidases act as catalysts for the physiological oxidation
of I- into I+ (i.e. HOI) (Vilter, 1995), Reaction (1), which can
then cross the plasma membrane. Apoplastic H2O2 is also consumed
for the oxidation of I- into I2. The oxidative formation of HOI in the
apoplast leads to a strong iodine solution in free diffusive contact with the
surrounding seawater. Upon oxidative stress, this iodine reservoir is
mobilized and a rapid, massive efflux of iodine occurs. Oxidation of iodide
results in the evolution of molecular iodine and volatile halogenated
compounds. In other words, HOI forms I2 via further reaction with
I-, as shown in Reaction (2), until equilibrium (2, -2) is achieved
(e.g. Lobban et al., 1985; Küpper et al., 1998; Hill and Manley, 2009).
I-+H2O2↔HOI+OH-
HOI+I-+H+↔I2+H2O
There have been a number of laboratory studies reporting that algae releases
organic and inorganic iodine after light-, chemical- and oxidation-induced
stress (e.g. McFiggans et al., 2004, and references therein; Palmer et al.,
2005; Hill and Manley, 2009).
During the springtime, solar radiation can penetrate through the relatively
thin Southern Ocean's sea-ice layer (see below) and sea-ice fractures,
reaching the micro-algae colonies that populate underneath and within
sea ice. In order to account for a diurnal pattern in the light-induced
iodine emissions from marine algae (Hill and Manley, 2009), the model includes
a parameterization of the iodine flux from the algae colonies following the
diurnal variation in actinic flux. We initialize the model using a biological
pre-concentration of 10-4 M iodide (micro-algae; Hill and Manley,
2009). We conducted sensitivity simulations, and these indicate that the release
of iodine is not significantly sensitive to a pre-concentration of 10-3
M ≤ [I-] ≤ 10-4 M.
The brine channels within sea ice and the BL at the sea-ice–air
interface. Following solar radiation incidence upon the algae colonies and
subsequent light-induced stress, intracellular iodine, equilibrated between
HOI, I- and I2, effluxes into brine channels of sea ice, and then
diffuses up to the BL, where it accumulates. The upward diffusion through the
brine channels is driven by an iodine concentration gradient. This gradient
arises from the concentration difference between the iodine emission point
(e.g. algae colonies) and the iodine content in the BL, which is
∼ 10-7 to 10-8 M (Chance et al., 2010; Atkinson et al.,
2012). For example, the upper limit for the concentration gradient can be up
to ∼ 10-4 M between the BL ([I-] ∼ 10-7 M) and the
iodine source in the algae colonies ([I-] up to 10-3 M).
Note that algae populations colonize the brine channel surfaces well into
the interior of the sea ice, close to the top of the sea-ice layer (Thomas
and Dieckmann, 2003). Hence, due to the occurrence of sea-ice fractures,
following springtime warming, it is also likely that, during the process of
sea-ice thinning and breakage, the algae colonies in the upper part of the
sea-ice layer will only be covered by a thin water film or be directly
exposed to air.
The comparatively thin Antarctic sea ice (mean sea-ice thickness
∼ 50 cm) (Thomas and Dieckmann, 2003) allows for the relatively fast
diffusion (see below) of iodine through sea-ice brine channels and further
release of I2(g), in addition to other iodine species,
such as IBr(g) and ICl(g), from the BL to the atmosphere. In the model we use
Fick's laws of diffusion, and diffusion coefficients for snow/ice (Shaw et
al., 2011; Loose et al., 2011; Callaghan et al., 1999; Mercier et al., 2005)
to compute the strength of the iodine flux, J, as a function of iodine
concentration gradient variability with time. Iodine fluxes are then obtained
by incorporating D (from 10-4 to 10-7 cm2 s-1) into
Fick's first law of diffusion (Shaw et al., 2011; Loose et al., 2011). Hence,
for the typical Antarctic sea-ice thickness (50 cm) we calculate
depth-dependent diffusion timescales (Table 1). Relevant diffusion timescales
range from 52 days at 30 cm (D=5× 10-5 cm2 s-1) days to 2.4 h at 2.5 cm (D= 1.3 × 10-4 cm2 s-1) (Table 1). According to
Eqs. (8) and (9), the iodine flux will decrease with time as the iodine
concentration gradient decreases due to accumulation of iodine in the BL.
Note that Loose et al. (2011), Callaghan et al. (1999) and Mercier et
al. (2005) show that the diffusion coefficients in Antarctic brine channels
range from 10-9 to 10-5 cm2 s-1 (fast diffusion
component), and the gas-phase diffusion coefficients in Antarctic sea ice
range from 10-7 to 10-4 cm2 s-1. Table 1 clearly shows
that, regardless of whether diffusion is
occurring through brine channels or via gas-phase diffusion, the timescale
will be fast. If the diffusion coefficient is between 10-7 and
10-9 cm2 s-1, iodine production would have to occur very
close to the surface to be relevant for polar springtime chemistry.
Diffusion timescale as a function of depth below sea-ice surface and
diffusion coefficient D via the diffusion length equation: L2= 4Dt, where, L is length (cm) and t is time (s).
We estimate iodine loss via measured loss rates of volatile organic iodinated
compounds (VOICs – CH3I, C2H5I, and C3H7I) assuming a VOIC concentration of ∼ 10-5 M (Shaw
et al., 2011). Maximum loss rates for each of these compounds were
∼ 2.1 nM h-1 (∼ 583 fmol s-1). Given the
range of iodine measured in algae (i.e. 10-7 to 10-3 M) and under
a unlikely scenario (as iodine is replenished within micro-algae) of no
iodine replenishment in micro-algae in conjunction with estimated VOIC loss
rates: (1) at 10-7, 10-5, and 10-3 M initial micro-algae
iodine concentration, iodine in the form of VOICs would be depleted in
∼ 4, 417, and 42 000 days, respectively. This unlikely baseline
scenario shows that the loss of iodine in the form of VOICs will not affect
the concentration of iodine emitted from micro-algae, especially at initial
iodine micro-algae concentrations above 10-5 M. Additional losses of
I2 and HOI via reaction with organic compounds are discussed in Sect. 6.
Therefore, even though our conservative model conditions show to be
sufficient for significant iodine release, it is possible that the iodine
concentration gradient between algae colonies and BL is larger than that used
in this study. For example, Hill and Manley (2009) showed that polar diatoms
were especially active in releasing iodine species – specifically, releasing
0.02–2.7 µmol HOI [mg total chlorophyll]-1 h-1, at
100 µmol L-1 iodide concentration. Note that
100 µmol L-1 iodide is 1 order of magnitude larger than the
iodide concentration used here in the CON-AIR model. One factor that can
influence the concentration gradient is the vertical extent of the algae
populations within the brine channels. For instance, the closer the algae
colonies are to the sea-ice surface, the lesser the possible losses of iodine
from reaction with organics in seawater during diffusion through brine
channels and the shorter the diffusion timescale for iodine to reach the BL
(Table 1).
Iodine exchange from Antarctic sea ice to the atmospheric boundary
layer. (a) Gas-phase I, I2, and IBr.
(b) Aqueous I-, I2, and IBr in the
BL. The post-sunrise pulse of I, shown in greater detail in the insert of
Fig. 2a, is the result of the night-time build-up of atmospheric I2.
Note that the emission of these species into the gas phase at night is slower
than their production rates in the condensed phase. With time, the I-
concentration becomes comparable to that of Br- and the reaction of HOI
with I-, forming I2, and competes with HOBr + Br- (forming
Br2) . The differences in the I2(aq) and
IBr(aq) concentration profiles are due to the fact that
the reverse rate constant to form HOI + Br-, starting from
IBr(aq), is orders of magnitude faster than that for
formation of HOI + I-, starting from I2(aq) (see
Supplement). Note that this figure corresponds to a diffusion timescale
∼ 6 days (D= 5 × 10-5 cm2 s-1 at
∼ 10 cm and D= 1.3 × 10-4 cm2 s-1 at
∼ 12.5 cm).
Model simulations and discussion
The model is initialized in October at local midnight at 75∘ S in
the Southern Hemisphere springtime. Figure 3 shows simulations of iodine
exchange between the BL and the atmosphere as a function of time. This figure
considers a diffusion timescale of 6 days (D=5× 10-5 cm2 s-1 at ∼ 10 cm and D= 1.3 × 10-4 cm2 s-1 at ∼ 12.5 cm) to
release enough iodine precursors to reach the IO levels (i.e. up to 20 pptv)
observed in coastal Antarctica (Saiz-Lopez et al., 2007a; Schönhardt et
al., 2008, 2012; Atkinson et al., 2012). Note, however, that the IO
concentration peak measured during a year-round campaign at Halley Bay
station occurred on 21 October (Saiz-Lopez et al., 2007a), which is about 70
days after spring sunrise at coastal Antarctica. The simulations in Fig. 3a
show that the nocturnal gas-phase I2 can reach concentrations of
7 × 108 molecule cm-3 over the course of 6 days, whereas
daytime I2 concentrations are much smaller due to its fast rate of
photolysis to form I atoms (Saiz-Lopez et al., 2004). Figure 3b shows that
the predicted I- concentration in the BL increases by 2 orders of
magnitude after 6 days of simulation due to the upward flux of iodine from
the algal colonies and the accumulation in the BL. I2(aq)
increases at a similar rate to I- since it forms primarily from
Reaction (2). Following Eq. (7), for a [I2(aq)]
∼ 2 × 10-8 M we estimate a transfer rate of I2 from
the BL to the gas phase of ∼ 1.5 × 105
molecules cm-3 s-1. The concentration of
IBr(aq), which forms from the reaction of Br- with
HOI, also increases in the BL in step with the increase of HOI.
Model simulations were run with a BL pH value of 8, similar to ocean water.
We have also assumed acidification of the BL and model runs for pH 4 have
shown a small enhancement in the release of gas-phase I2. As the model
simulation evolves with time, the strength of the iodine flux from the
phytoplankton and the iodine concentration increase in the BL will be the
major drivers determining the timescale as well as the concentrations of
photo-labile reactive iodine precursors released from sea ice.
Also note that, as the Antarctic springtime progresses, there are two factors
enhancing the accumulation of I- in the BL: (i) the phytoplankton bloom
associated with high iodine emissions and increase in solar irradiance
reaching the sea-ice surface and (ii) the thinning of the sea ice and more
frequent occurrence of brine channels favouring faster upward transport
through the ice, as well as the break-up of sea ice, which exposes phytoplankton
colonies, and their associated iodine emission, directly to the atmosphere.
Modelled concentrations of gas-phase iodine species
resulting from the emission of I2 from sea ice over 6 days from the
start of the simulation. Following the build-up of I2 during the
preceding night, the model predicts a post-sunrise pulse of IO followed by a
diurnal cycle shaped by solar radiation. The HOI (from IO + HO2) and
HI (from I + HO2) profiles track the diurnal cycle of HO2 and
solar radiation. By contrast, IONO2 will photolyse very efficiently
during the day, yielding its typical diurnal cycle with maxima in mid-morning
and late afternoon.
Figure 4 shows an example of the gas-phase chemistry resulting from I2
release to the atmosphere, following the model run conditions shown in
Fig. 3. Atomic iodine reacts with atmospheric O3 to form IO, and this
radical then self-reacts to yield OIO. The calculated concentrations of IO
can reach 2 × 108 molecules cm-3; these levels are in
good agreement with average boundary layer concentrations of the molecule
recently measured at coastal Antarctica both from the ground (Saiz-Lopez et
al., 2007a; Atkinson et al., 2012) and from satellite platforms
(Schönhardt et al., 2008, 2012). The computed O3, also plotted in
Fig. 4a, shows a substantial rate of depletion due to iodine chemistry of
0.25 ppb h-1. This is almost twice as fast as that calculated from
bromine-mediated chemistry alone (0.14 ppb h-1 for typical Antarctic
springtime boundary layer BrO mixing ratios of 10 pptv). Figure 4b shows the
diurnal profiles of gas-phase HOI, HI and IONO2. These three species can
be deposited back from the gas phase onto the sea-ice surface and
subsequently converted to aqueous HOI. The set of reactions involved is
summarized as follows (iodine species are in the gas phase unless indicated):
HOI→HOI(aq),HOI(aq)+I-+H+→I2(aq)+H2O,I2(aq)→I2,I2+hν→2I,2(I+O3)→2(IO+O2),2(IO+HO2)→2HOI+2O2.
Another point to consider is that this mechanism potentially establishes a
synergy between the biologically induced emissions of iodine and the trigger
of bromine release from Antarctic sea ice. The model results show that the
increase in iodine content in the BL will also trigger the catalytic release
of bromine from sea ice via formation and subsequent release of IBr to the
gas phase (see Fig. 3), which, following photolysis, will provide a source of
reactive bromine in the Antarctic atmosphere.
We also propose that similar to the inorganic iodine release mechanism,
algal emissions of iodocarbons followed by transport and accumulation in the
top of the sea-ice layer may arise in phase equilibration of organic iodine
from ice-covered ocean areas to the atmosphere.
Finally, we suggest that this mechanism is more efficient for the Antarctic
sea-ice environment than for the Arctic due to physical constraints such as
greater mean sea-ice thickness (e.g. ∼ 3 m) and smaller algal
population in the Arctic (Thomas and Dieckmann, 2003). Due to the
non-linearity in the system (see Eqs. 8 and 9), our calculations show that
the diffusion timescale of iodine species through Arctic sea ice is
∼ 40 times slower than that in the Antarctic. This is an upper limit
for Arctic iodine emissions through the proposed mechanism since algal
colonies are less predominant in the Arctic than in the Antarctic and
propagation of solar irradiance through ice will also be largely limited due
to thicker sea ice. This will greatly limit the overall metabolic production
of iodine species. However, it cannot be ruled out that, when the Arctic sea
ice melts, the phytoplankton colonies will be directly exposed to air and
therefore constitute a potential source of iodine in the Arctic atmosphere.
The difference in Arctic and Antarctic sea-ice microstructure – that is,
predominantly columnar ice in the Arctic versus frazil and platelet ice in Antarctica may also influence the observed differences in halogen release over sea ice.
Uncertainties and future work
Although it is beyond the scope of this manuscript, below we discuss
additional loss and production pathways for iodine and VOICs within sea ice
(Table 2) which should be addressed in future model studies.
Reactions of I2 and HOI with dissolved organic matter (DOM) in sea ice/snow. Assuming I2 concentrations
of ∼ 10-7, 10-5 and 10-3 M and a first-order loss rate ∼ 7 × 10-3 s-1
(for coastal water) and 5 × 10-5 s-1 (for open water) of I2 with dissolved organic matter (DOM)
(Truesdale, 1995; Carpenter et al., 2013), the lifetime (without replenishment) of I2 would be ∼ 2.4 and ∼ 333
minutes, respectively. We will also investigate the equilibrium reaction of I2+ I-↔ I3-
(O'Driscoll, 2008), whose equilibrium lies well to the right, forming the trihalide ion; still, its forward reaction rate is 3
orders of magnitude slower than the primary reaction releasing I2 – that is, HOI + I-+ H+↔ I2+ H2O.
Abiotic release of iodine from water surfaces; for instance, Chance et al. (2010) observed surface seawater iodide levels
of up to 150 nM in summer off the coastal western Antarctic Peninsula. Heterogeneous reaction with 30 ppb atmospheric ozone could
release ∼ 2 × 107 molecules cm-2 s-1 for I2 and 3.5 × 108 molecules cm-2 s-1
for HOI from O3+ I- (Carpenter et al., 2013).
Via haloperoxidase activity, biogenic emissions of iodine are another viable release mechanism. Under optimum assay conditions
using Porosira glacialis, a centric diatom, Hill and Manley (2009) observed release rates up to 271 fmol HOI cell-1 h-1.
Estimated production rates are highly dependent on the amount of biomass, as well as I- and H2O2 concentrations.
Much lower activity was observed at iodide concentrations closer to natural seawater. Based on data shown in Hill and Manley (2009),
release rates may be a factor of 100 lower under ambient H2O2 and iodide conditions, e.g. ∼ 0.03 μmol HOI [mg total Chl]-1 h-1. Using representative chlorophyll concentrations from Sturges et al. (1997) yields a very high
production rate of ∼ 1 × 1010 molecules cm-2 s-1 HOI. The majority of the biomass producing this
will be at the base of the ice; therefore losses of HOI will occur before release to the atmosphere.
Production of organic iodine within leads and polynyas and near sea ice. To explain the atmospheric levels of organoiodines
and IO at Hudson Bay, Mahajan et al. (2010) postulated CH2I2= 1 × 106,
CH2IBr = 2 × 109, CH2ICl = 5 × 107,
and CH3I = 2 × 108 molecule cm-2 s-1 from open leads 15 h upwind of measurements. At Hudson Bay,
Carpenter et al. (2005) measured high concentrations (i.e. ∼ 1–3 pptv) of reactive dihalomethanes. Yet, organic iodine
compounds are not ubiquitously high in polar regions (Carpenter et al. 2007; Atkinson et al., 2012; Granfors et al., 2015).
Nitrate and H2O2 photochemistry produces OH (Anastasio et al., 2007;
Dubowski et al., 2001; Chu and Anastasio, 2003), which may alter the pH of the snowpack. Given that halogen chemistry
is pH-dependent; such photochemical reactions may be interrelated through
this
context with halogen chemistry, although currently there is a lack of multi-solute experimental data to adequately simulate this process.
Still, O'Driscoll et al. (2006) report the production of trihalides via iodide- and nitrite-doped ice matrices.
Summary of additional iodine production and loss processes.
a Truesdale et al. (1995),
b Palmer et al. (1984),
c O'Driscoll et al. (2008),
d Carpenter et al. (2013),
e Hill and Manley (2009),
f Sturges et al. (1997),g Mahajan et al. (2010),
h Carpenter et al. (2005).
Summary and conclusions
A mechanism for iodine release from sea ice has been proposed. This mechanism
incorporates the coupling between stress-induced biological emissions of
iodine, diffusion of iodine through sea ice, and phase equilibration to the
atmosphere. In order to quantitatively investigate the feasibility of the
mechanism a multiphase chemical model has been developed. Model simulations
for the coastal Antarctic springtime show that the release of photo-labile
inorganic iodine (i.e. I2, IBr, ICl) could account for the observations
of elevated IO in this environment, which is primarily sourced near the
surface of sea ice (i.e. ≥ 30 cm for D= 5.5 × 10-5 cm2 s-1 and ≥ 50 cm at D= 1.3 × 10-4 cm2 s-1). Most likely, the overall
mechanism involves a combination of biological emissions of iodine
simultaneously at different depths within the sea-ice column. This process
may also trigger reactive bromine release from sea ice via gas-phase
equilibration and subsequent photolysis of IBr. In addition, following the
same mechanism, organic iodine may also be released from sea ice to the
atmosphere. Lastly, it appears that coastal Antarctic sea ice is not alone in
emitting iodine, as recent measurements have reported 3.4 ± 1.2 pptv IO
in the Arctic over open-water polynyas that form in the sea ice (Mahajan et
al., 2010). Also, iodine has been detected in growing particles over coastal
sea ice near Greenland (Allan et al., 2015). The smaller amounts of IO
measured over the Arctic may be indicative of differences in Arctic
vs. Antarctic sea-ice thickness, micro-algae amount, sea-ice
microstructure, salinity, porosity, and temperature. Therefore, this
mechanism may also govern the emission of iodine (including VOICs) measured
in the Arctic as well. We also acknowledge here that the efflux of HOI and
I- via diatoms/micro-algae likely forms I2 rapidly as a function of
depth, which could also be coupled to the diffusion of I- and HOI to the
top layers to be released at the surface; this possibility (given the rapid
forward reaction of HOI(aq)+ I-+ H+→ I2(aq)+ H2O) may also give
insight into lower I- concentrations measured in brine channels compared to its pre-concentration in algae.
The Supplement related to this article is available online at doi:10.5194/acp-15-9731-2015-supplement.
Acknowledgements
During the inception of this project, A. Saiz-Lopez and C. S. Boxe were
supported by appointments to the NASA Postdoctoral Program at the Jet
Propulsion Laboratory, administered by Oak Ridge Associated Universities
through a contract with NASA. Research at the Jet Propulsion Laboratory, California Institute of
Technology, was also supported by the NASA Upper Atmosphere Research and
Tropospheric Chemistry programs. L. J. Carpenter thanks the UK NERC (grants
NE/I028769/1 and NE/D006538/1) for funding. We thank Erik Campbell for
producing Fig. 1. We are grateful to John Plane, Stanley Sander,
Ross Salawitch and Rosie Chance for very helpful discussions and comments on
this work. Edited by: M. C. Facchini
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