Introduction
Iodine chemistry plays an essential role in the marine boundary layer (MBL)
due to its effect on the destruction of tropospheric ozone, perturbation of
the HOx–NOx cycle and the formation of new particles and cloud
condensation nuclei, thereby leading to changes in the global radiative
forcing (Hoffmann et al., 2001; von Glasow and Crutzen, 2003; O'Dowd and
Hoffmann, 2005; Bloss et al., 2005; Huang et al., 2010a, b). This essential
role of iodine and of other activated halogens is shown in field
measurements in the marine boundary layer (MBL), laboratory chamber
experiments or incubation experiments of different algae and in atmospheric
models (Carpenter, 2003; Küpper et al., 2008; Kundel et al., 2012;
McFiggans et al., 2000). The biogeochemical cycle of iodine is controlled by
large iodine exchanges from the oceans to the atmosphere, driven by marine
biotic and abiotic production (Schall et al., 1997). Volatilized species are
photolabile iodocarbons like CH2I2, CH3I, C2H5I,
CH2ICl, CH2IBr and molecular iodine (I2). Marine species like
macroalgae and microalgae play a dominant role in the emission of these
compounds (Carpenter et al., 1999; Huang et al., 2013; Saiz-Lopez and Plane,
2004).
Since molecular iodine and iodocarbons are photochemically unstable
(lifetimes between about some tens of seconds for I2 and a few days for
CH3I), they are photolysed under UV–visible light to form I⋅ atoms, which are then instantly oxidized by ozone to form the iodine
monoxide radical IO (g) (Hoffmann et al., 2001; Saiz-Lopez et al., 2006).
Further oxidation reactions of IO in the gas phase then can form low-volatility
iodine oxides (IxOy) which may nucleate under certain
conditions and form new particles.
Recently it was proposed that the ozone loss over the tropical Atlantic
Ocean was higher than calculated from global atmospheric models, and that
this additional ozone destruction is induced by halogens such as bromine and
iodine (Read et al., 2008). Biogenic emissions, such as the already-studied
iodocarbon emissions by phytoplankton species – e.g. coccolithophorids,
diatoms and chlorophytes (Colomb et al., 2008) – are too low to explain the
differences in model calculations and observations (Mahajan et al., 2010);
therefore additional sources of the reactive iodine species are discussed,
one of them being the surface reaction of ozone with seawater.
Garland and Curtis first discovered that the emission of molecular iodine
from the surface of artificial and natural seawater is proportional to the
ozone concentration at the air–water interface (Garland and Curtis, 1981).
Sakamoto and co-workers examined the reaction mechanism of the iodide
oxidation by ozone at the air–water interface, resulting in the formation of
the intermediates IOOO- and HOI and the emission products IO and
I2 (Sakamoto et al., 2009). Further laboratory experiments show that
different organics affect the reaction of iodide with ozone; e.g. fulvic
acid enhances the I2 formation but not the formation of IO (Hayase et
al., 2010, 2012).
Since the formation of I2 and IO from the air–water interface is
dependent on the iodide concentration in seawater, the reaction path found
by Garland and Curtis may explain elevated iodine emissions in areas of
higher phytoplankton activity (Garland and Curtis, 1981). The ability of
different phytoplankton, e.g diatoms, to reduce iodate, which is ubiquitous
in the open ocean, to iodide was shown for natural and elevated iodate
concentrations (Wong et al., 2002; Chance et al., 2007) and for the different
growth states (Bluhm et al., 2010) of the phytoplankton cultures. A
correlation of iodine species in the particle phase and average chlorophyll
exposure of air masses along back trajectories was found by Lai et al. (2011),
indicating the link between phytoplankton activity and emission of
atmospheric iodine.
Since the formation of I2 and IO is correlated with the iodide
concentration (Sakamoto et al., 2009) and the iodide concentration of surface
waters is correlated with phytoplankton (Bluhm et al., 2010), this study
investigates links between iodide concentrations in microalgae-containing
seawater and abiotic formation and emission of I2, utilizing laboratory
experiments of the reaction of the seawater surface with ozone.
Materials and methods
Experimental set-up
Two diatom cultures (M. helysia and Porosyra glacialis, both from the Alfred Wegener Institute/Sylt) were kept
in F/2 seawater medium for growing. These media were prepared from filtered
natural seawater from the shores of Sylt and additional nutrients which the
diatoms need to grow (0.88 mmol NO3-, 0.04 mmol PO43-
and 0.01 mmol SiO32-) and which is a common used medium as
described by Guillard et al. (1975) and Kraberg et al. (2012). Both cultures
were incubated in the F/2 medium at 16 ∘C with
12 h light–12 h dark cycling (LUMILUX Plus Eco daylight lamp;
ca. 40 µmol PAR) for 4 weeks prior the experiment. Just before the emission
experiment, the algal suspensions were diluted in a 2:1 ratio in F/2 medium
and homogenized by stirring. In addition to the diatom cultures, a plankton
concentrate was collected from the North Sea (55∘01.562′ N,
8∘27.113′ E) on 24 May 2012 using a 80 µm and
200 µm Apstein plankton net and diluted using the same medium as for the
diatom cultures. Microscopic observations showed that the plankton
concentrate sample was dominated by colonies of the haptophyte Phaeocystis sp. and only a
low amount of diatoms were present.
Experimental set-up of the chamber with the phytoplankton
suspension.
For each experiment 1.5 L of the sample (i.e. diatom suspension, natural
plankton concentrate or background (F/2 medium)) was introduced into a glass
chamber tube (10 L), shown in Fig. 1, and three
magnetic stirrers were switched on immediately. A continuous flow of
synthetic air (3.4 L min-1) was channelled over the stirred algae
suspension (surface area 2250 cm2) in the first experiment with no
ozone and in the second experiment with elevated ozone levels of 100 ppb.
The ozone was generated using an UV radiation source, and the resulting ozone
levels were measured using an ozone analyser (Dasibi Environmental Corp.
Model 1008-RS, Glendale, USA). To measure the emission of I2 and
halocarbons, α-cyclodextrin-coated denuders (Huang and
Hoffmann, 2009; Huang et al., 2010c) and adsorption tubes (Kundel et al., 2012)
were mounted at the other end of the tube chamber together with the
ozone monitor. The chamber outflow was sampled using two membrane pumps, one
with 0.50 L min-1 for the denuders and the other using 0.15 L min-1
for the adsorption tubes. To assure an overpressure over the
sampling time, a U-shaped tube filled with ultrapure water was mounted in
the centre exit of the glass chamber to measure the overpressure
hydrostatically. The whole set-up was wrapped with aluminium foil to prevent
photolysis of I2 and halocarbon compounds. Potential wall losses of
I2 and halocarbons were investigated using diffusion (I2) and
permeation (halocarbons) test gas sources with a dilution chamber; no wall
losses were observed within the precision of the measurements for the three
concentrations tested – 70 ppt, 270 ppt and 700 ppt for iodine and 0.2 ppt,
2.4 ppt and 8.3 ppt for the halocarbons – within the stated gas
flows.
To monitor the emissions of I2 and halocarbons from the liquid samples,
an evaporation standard was added to the microalgal suspension in order to
highlight any problems related to air sampling. This standard was
1,3-dibromopropane diluted in ultrapure water (500 µL of
0.94 µg L-1, which was then diluted with the sample to 1.5 L). The standard
was chosen given the results from a first set of experiments with M. helysia and
Coscinodiscus wailesii which show no detectable traces of this compound. We decided not to add any
iodine-containing compounds to prevent interferences with the I2
emission.
Halocarbon measurements
Air samples of 6.75 L sampling volume were pre-concentrated at a flow rate
of 150 mL min-1 on thermal desorption tubes filled with 100 mg Tenax TA
60/80 and 150 mg Carbotrap™ 20/40, both provided by Supelco
(Bellefonte, PA, USA). The samples were analysed using a self-made thermal
desorption device mounted on a gas chromatograph (GC, TraceGC, Thermo
Scientific, Dreieich Germany) and mass spectrometer (MS, PolarisQ, Thermo
Scientific, Dreieich, Germany). During the desorption period of 6 min the cryotrap was cooled to -160 ∘C. Afterwards the cryotrap was
rapidly heated to 270 ∘C for injection. The analytes were
separated on a DB624 Durabond column (60 m; 0,32 mm; 1,8 µm FT) using
helium as carrier gas with a constant pre-column pressure of 0.5 bar. The
temperature program was 55 ∘C (4 min), ramp with 5 ∘C min-1 to 120 ∘C (4 min) and ramp with
8 ∘C min-1 to 200 ∘C (4 min). Halocarbons were detected using a
mass spectrometer in negative chemical ionization (NCI) mode with methane as reagent gas (2.5 mL min-1), the primary electron energy was set to 120 eV and an emission
current of 50 mA in single ion monitoring mode (SIM) was used. Iodinated
compounds (CH3I, C2H5I,
CH2ICl, CH2IBr, CH2I2, 1-C3H7I,
2-C3H7I, 1-nC4H9I, 2-nC4H9I,
1-iso-C4H9I) were quantified using m/z 127, and brominated
compounds (CH2Br2, CH3Br, 1,3-C3H6Br2) were quantified using m/z 79 and 81 at a 1:1 ratio.
A five-point calibration was done in the range between 0.01 ng and 1 ng
using the continuously diluted output of a permeation test gas source
(Thorenz et al., 2012). The detection limits for the individual iodocarbons
were 0.003–0.088 ppt and for the bromocarbons were 0.004–0.009 ppt. For
each series of measurements, the calibration was done in triplicate
(precision of method: 3–13 %).
I2, iodide and iodate measurements
Sampling of gaseous I2 was performed using the denuder technique
described by Huang and Hoffmann (2009). Brown glass denuder tubes (6 mm inside diameter,
50 cm length) were coated using an α-cyclodextrin suspension
(2.5 mg mL-1 in methanol) and sealed with polypropylene caps. Before sampling
the denuders were stored in a refrigerator. For sampling the denuders were mounted
vertically with a glass tube of 15 cm upstream to achieve laminar flow. The
sampling flow was 500 mL min-1 for 45 min. After sampling the denuders
were sealed and stored in a refrigerator until derivatization. For derivatization
the α-cyclodextrine coating was eluted with ultrapure water (20 mL);
then 25 µL N,N-dimethylaniline (1 µg mL-1 in methanol),
500 µL phosphate buffer (pH 6.4) and 500 µL 2-iodosobenzoate
(4 mg mL-1) were added, and the mixture was shaken for 2 h. After
adding 3 mL sodium acetate, the sample was extracted with 100 µL cyclohexane and 100 µL 2,4,6-tribromoaniline (internal standard: IS)
in cyclohexane (250 ppb).
Iodide and iodate were derivatized from seawater to form the same product as
described for I2. Iodide was oxidized to form I2 by using
iodosobenzoate, and iodate was reduced first to iodide and then oxidized to
form I2. Ten-millilitre aliquots of seawater were analysed for iodide and for
total iodine; iodate was calculated by difference. The method for iodide
derivatization was slightly changed from the one described by Mishra et al. (2000).
The use of sodium hydrogen sulfite as an agent to reduce iodate to
iodide is described by Schwehr and Santschi (2003).
Iseawater=I-+IO3-
To measure iodide, 10 mL seawater was mixed with 1 mL ethylenediaminetetraacetic acid solution (0.5 %), 500 µL phosphate
buffer, 500 µL N,N-dimethylaniline and 500 µL iodosobenzoate and
shaken. After adding 3 mL sodium acetate the sample was extracted with
100 µL cyclohexane and 100 µL 2,4,6-tribromoaniline (IS) in
cyclohexane (250 ppb).
To measure iodate an aliquot of 10 mL seawater was mixed with 1 mL ethylenediaminetetraacetic acid solution (0.5 %), 1 mL hydrochloric acid
(3.7 %) and 500 µL sodium hydrogen sulfite solution
(283.9 µmol L-1) to reduce the iodate. Afterwards 500 µL sodium acetate,
4 mL phosphate buffer, 500 µL N,N-dimethylaniline and 500 µL iodosobenzoate were added. After shaking the sample was again extracted with
100 µL cyclohexane and 100 µL 2,4,6-tribromoaniline (IS) in
cyclohexane (250 ppb).
One microlitre of the cyclohexane extract was injected to the GC–MS system
(6850 GC & 5973 MS, Agilent Technologies, Waldbronn, Germany) at a
constant flow of 1 mL min-1 of helium (99.999 %), and the
chromatographic separation was performed using a capillary column FS Supreme
5 MS with a length of 30 m, inner diameter of 0.25 mm and film thickness of
0.25 µm (CS Chromatographie Servieve, Langenwehe, Germany) with a
temperature program starting at 50 ∘C (for 3 min), then heating up
at 30 ∘C min-1 to 220 ∘C (for 3 min). The mass
spectrometer measured in electron ionization mode at 70 eV; the specific
fragments of the product 4-iodo-N,N-dimethylaniline was extracted at m/z 247
(M+) and of the internal standard 2,4,6-tribromoaniline at m/z 329 (M+).
Chlorophyll measurements
The analytical method for chlorophyll a (Chl a) measurements
is described by Edler et al. (1979). An aliquot of 50–100 mL water samples were
filtered on glass fibre filters (Whatman, Grade GF/F). The dry filters were put in
polypropylene vials and extracted with 7.5 mL acetone. The extract was
stored together with the filter in a dark refrigerator at 3 ∘C overnight
and centrifuged the next day (5500 rpm, 7 min) at 5 ∘C. The
absorption of the supernatant was measured against acetone using an Uvikon
XL double-beam spectrophotometer at λ=750, 663, 645
and 630 nm. To calculate the concentration of Chl a the equation of
Jeffrey and Humphrey (1975) was used. Chl a can be a good indicator
for microalgae biomass (Roy 2010; Bluhm et al., 2010; Colomb et al., 2008)
and has been used to calculated emission rates of iodine-containing
volatiles from phytoplankton. This calculation was not used here, since the
mechanisms of synthesis and release of these iodine-containing gases is
still unclear. All gaseous compounds in this study are therefore given as
measured mixing ratio, and the Chl a value of the corresponding algae
suspension is also given.
Halocarbon emission rates, concentrations of chlorophyll α,
iodide and iodate in the four different sample suspensions.
Sample
F/2 medium background
P. glacialis
M. helysia
Plankton concentrate
Range (Mean)
Range (Mean)
Range (Mean)
Range (Mean)
CH3I
pmol min-1 m-2
0.17–0.72 (0.35)
0.21–0.69 (0.45)
0.32–0.82 (0.53)
0.08–0.37 (0.19)
CH2ICl
pmol min-1 m-2
0.02–0.22 (0.11)
0.02–0.22 (0.16)
0.04–0.22 (0.18)
0.02–0.12 (0.07)
CH2I2
pmol min-1 m-2
0.02–0.08 (0.07)
0.27–0.44 (0,36)
0.21–0.50 (0.37)
0.04–0.09 (0.07)
CHBr3
pmol min-1 m-2
1.76–1.90 (1.81)
1.99–2.17 (2.09)
1.75–2.17 (2.09)
1.75–2.33 (1.82)
Chl a
µg L-1
n.d.
257
927
2.5
Iodide
nmol L-1
6.6–15.6 (10.4)
7.3–19.7 (12.7)
9.9–21.9 (16.8)
3.5–9.5 (6.5)
Iodate
nmol L-1
402–538 (428)
408–478 (448)
397–499 (446)
424–478 (442)
1,3-C3H6Br2∗
pmol min-1 m-2
7.77 ± 0.04
7.78 ± 0.59
7.77 ± 0.99
7.69 ± 0.07
ΣIodocarbon/Chla
pmol/g
n.d.
19.75
6.06
694.88
∗ Evaporation standard given as mean ± standard deviation.Chl a was measured for each sample once halocarbons, iodide, iodate mean
values and ranges had been calculated from six replicates
ΣIodocarbon/Chla. Iodocarbon emissions were summed for
the eperimetntal conditions (time and surface area) and normalized to Chl a in the watery phase.
Results and discussion
Halocarbons
The emission rates of the natural halocarbons and the evaporation standard,
given in Table 1, were calculated by the amount measured in the adsorption
tubes divided by the emission time and the surface area of the suspension
sample (pmol min-1 m-2). The halocarbon emission rates showed no
effect on the different ozone levels; therefore the data for each sample are
summarized for high- and low-ozone conditions. An evaporation standard was
added to the different samples to recognize differences in emission rates of
the organic compounds from the aqueous phase. The standard was added in a 10-
to 100-fold excess compared to natural concentrations of bromocarbons in
Atlantic seawater (Carpenter et al., 2000) to reduce the effect of natural
1,3-dibromopropane, which may alter the mixing ratio of the evaporation
standard measured. In the chosen concentration a natural abundance would
change the result only by 1–10 % compared to the spike solution. The
results of the measurements of 1,3-dibromopropane showed very constant
values, as can be seen from the low standard deviation between the different
samples and replicates. This result indicates a stable and reliable
experimental set-up in terms of evaporation of volatile compounds from the
water surface and of the mixing of the bulk water.
The measured emission rates of the natural halocarbons show that the
brominated compound, CHBr3, is elevated compared to the iodocarbons
emission rates. This result fits to observations of the natural abundance of
halocarbons in seawater as described in earlier studies (Roy et al., 2011).
The emission rate of CHBr3 is higher for the two diatom cultures (M. helysia and
P. glacialis) than for the plankton samples containing Phaeocystis sp. and the background. Again, this
result matches field and laboratory data showing a link between elevated
CHBr3 concentrations in seawater and the simultaneous occurrence of
diatoms (Colomb et al., 2008; Quack et al., 2007; Moore et al. 1996).
The iodocarbon emissions in experiments using the background (F/2 medium)
and the plankton concentrate were dominated by CH3I, followed by
CH2ICl and CH2I2. However, for the diatom cultures,
CH2I2 was the dominant iodocarbon emitted, with CH3I and
CH2ICl both showing lower emission rates. The emission of iodocarbons
from the F/2 background is not surprising for two reasons; first, the medium
was produced from natural shoreline filtered water, which already may
contain iodocarbons (Wong and Cheng, 1998). The second reason may be related
to iodocarbon-producing bacteria (Amachi et al., 2001, 2003).
These bacteria could have been present and active in the natural seawater
used to produce the F/2 medium, since it was not sterilized prior to
use. Additionally, the emission rates of CH2I2CH2ICl and
CH3I in the diatom samples P. glacialis and M. helysia were significant higher compared to the
background (Wilcoxon rank sum test p=0.00032 and p=0.00007,
respectively). This increase in emission can be explained by the capability
of the diatoms to produce iodocarbons, which had already been reported by
Moore et al. (1996). To compare the natural plankton concentrate with the
cultivated diatom cultures and the background, one must keep in mind that Chl a
concentrations are biomass tracers reflecting the abundance of
phytoplankton. The results for the Chl a measurement, given in Table
1, clearly show that the natural plankton concentrate contains less biomass
than the cultured diatoms. Therefore, we conclude that the lower iodocarbon
emissions of the plankton concentrate compared to the diatom cultures is
partly due to lower biomass density. The lower iodocarbon emission rates in
the natural plankton concentrate could also be related to iodine uptake of
naturally occurring microalgae (van Bergeijk et al., 2013). The emission flux
summed for the three iodocarbons in the four samples' background, plankton
concentrate, P. glacialis and M. helysia, was in the range 0.21–1.02,
0.14–0.58, 0.50–1.35 and 0.57–1,53 pmol min-1 m-2, respectively. We are not aware of
emission studies investigating the flux of iodocarbons from microalgae
suspensions to directly compare these results. To establish a connection to
other experimental observations, the results listed above are compared to
incubation studies of marine aggregates producing iodocarbons and calculated
emission fluxes in coastal, seaweed-rich regions. Hughes et al. (2008)
measured the iodocarbon production of different marine aggregates to be
within 0.71 to 6.90 pmol h-1 L-1. The production rate is difficult
to compare to the presented results, since the flux in our study is based on
the production by the microalgae species and evaporation from the surface,
whereas Hughes et al. (2008) measured the production in the aqueous phase.
Jones et al. (2009) calculated iodocarbon emissions at a sampling site
surrounded by fields of macro algae in open-sea water at Roscoff, France.
The flux of iodocarbons was estimated to 85.28 pmol min-1 m-2, 2
orders of magnitude higher than the flux obtained in the present study. Thus
it appears that on an areal basis the natural populations of microalgae
studied here are much less prevalent emitters of iodocarbons than seaweeds
and marine aggregates.
Iodide and iodate
The concentrations of iodide and iodate in the different samples are shown
in Table 1. For each sample, the mean and range for six replicates are
shown; no differences in iodide and iodate concentrations were observed
under elevated- (100 ppb O3) and low-ozone (0 ppb O3) conditions.
The iodate concentrations in the background and in the three plankton
samples were in the same range, with mean concentrations between 438 and
448 nmol L-1. These iodate concentrations are in the range measured for the
open ocean of 400 to 500 nmol L-1 iodate in most oceanic regions (Bluhm
et al., 2011). The ubiquity of iodate suggests that its concentration is not
a limiting factor.
The iodide concentrations in the two diatom cultures, P. glacialis and M. helysia, are slightly
elevated, with mean values of 12.70 nmol L-1 and 16.84 nmol L-1,
respectively, compared to the background iodide concentration of 10.35 nmol L-1 and the
plankton concentrate iodide concentration of 6.47 nmol L-1. This enhanced iodide concentration indicates the reduction of
iodate by the two diatom cultures, which was also found by Bluhm et al. (2010)
and Wong et al. (2002) for different phytoplankton species. Such a reduction
of iodate to iodide will result in a decrease in the iodate concentration;
however, for the measured iodate concentration in this study the expected
decrease falls within the analytical precision of the measurement. The
iodide concentrations in all samples are comparable with oceanic surface
water concentrations, for example around 10–30 nmol L-1 in the Weddel
Sea surface water (Bluhm et al., 2011).
The low iodide concentration of the plankton concentrate sample compared to
the background sample is surprising but may be assigned to an overall low
level of different nutrients, like phosphate and silicate, in the Wadden Sea
of Sylt during springtime (Weisse et al., 1986), although the level of iodate was
consistent. Another possible reason for the low iodide concentration in the
plankton concentrate could be iodine uptake by microalgae present in the
natural plankton sample (van Bergeijk et al., 2013).
Ozone measurements
The results of the ozone measurement for the samples' background, P. glacialis, M. helysia and the
plankton concentrate were normalized against a background measurement
obtained using ultrapure water in the chamber. This was performed in order
to account for losses of ozone through wall reactions, losses on the water
surface and losses due to droplet formation from stirring. The ozone
consumption was calculated using a continuous stirred-tank reactor (CSTR)
approach with 668 ng min-1 ozone (100 ppb) introduced into the chamber
(total volume: 10 L; flow: 3.4 L min-1; and residence time: 2.94 min).
The difference between the introduced ozone flow and measured ozone flow is
considered as consumed ozone, due to the oxidation of iodide and other
ozone-depleting reactions in the samples. To calculate the consumed ozone, the
flow rate was summarized over 45 min of the experiment. Ozone consumption
was clearly observed for all samples. The background sample showed the
weakest ozone consumption of 58 nmol, followed by the sample of P. glacialis with
186 nmol and the plankton concentrate with 253 nmol. The highest ozone
consumption was shown by M. helysia with 335 nmol.
Iodine emission rates normalized for the surface area of the
different samples at 0 and 100 ppb ozone. The error bars represent the
standard deviation of the three replicates of each experiment.
I2 emissions
The I2 emission rate was calculated by dividing the amount of I2
by the sampling time and the suspension surface area. The results for the
four samples are shown in Fig. 2. The background and the two diatom
samples, M. helysia and P. glacialis, show significantly higher emission rates when the ozone level
is elevated (100 ppb O3) compared to conditions where no ozone is
present (0 ppb O3). The difference between the high- and low-ozone
conditions is small for the background, increases for the P. glacialis sample and is
highest for the M. helysia sample. The plankton concentrate does not show a
significant dependence of the I2 emission rate on the ozone level. The
ozone-dependent increase in the I2 emission rate of the other samples
indicates that iodide, which is present at the air–water interface, is
oxidized by ozone to form I2, which is consistent with the
results from artificial and natural seawater (Garland and Curtis, 1981;
Sakamoto et al., 2009).
Correlation of the change in the I2 emission and the iodide
concentration in the microalgae suspension.
Figure 3 shows the change in I2 emission rate
([I2 at
100 ppb ozone]-[I2 at 0 ppb ozone]) of the different
samples as a function of the iodide concentration measured in the bulk
water. A linear correlation fits the data well with a Pearson's coefficient of
R2=0.998. This behaviour indicates a direct proportional
relationship, which was also seen by Sakamoto et al. (2009) for small iodide
concentrations (0–5 mmol L-1). Carpenter et al. (2013) also observed
that the I2 emission is dependent on the aqueous iodide concentration.
The proposed reaction sequence, as shown in Eqs. (1–5), explains the
relationship between the iodide concentration in the aqueous phase and the
I2 emissions (Sakamoto et al., 2009).
I-(aq)+O3(g or interface)→IOOO-(interface)IOOO-(interface)→IO-(aq)+O2(aq)IO-(aq)+H+↔HOI(aq)HOI(aq)+I-(aq)+H+↔I2(aq)+H2OI2(aq)→I2(g)
The plankton sample does not show an elevated I2 emission at 100 ppb
ozone compared to zero ozone. This observation indicates that in the
plankton sample an additional I2 loss process takes place. Reactions or
partitioning of I2in an organic surface layer, which was discussed in
Carpenter et al. (2013), would be one possibility to explain these results.
In fact the specific microalga found in the plankton concentrate, Phaeocystis sp., is known
to produce high amounts of organic matter (Eberlein et al., 1985). An
alternative explanation is the low iodide concentration in the plankton
concentrate, which may be related to iodide uptake by the naturally occurring
plankton communities. The iodide concentrations and ozone mixing ratios in
this study represent more likely natural conditions compared to the study of
Sakamoto et al. (iodide concentration from 0.01 to 50 mmol L-1 and
ozone mixing ratio from 2 to 298 ppm). However, the results presented here
demonstrate that, even under low iodide concentrations, representative of
natural conditions of the MBL, a significant formation of I2 by the
ozone driven oxidation of iodide at the air–water interface takes place,
until the iodide concentration gets too low.
Comparing the I2 and iodocarbon emission rates, it is clear that the
volatile iodine emissions are dominated by I2 . Therefore I2
emissions from natural seawater surfaces are more relevant for atmospheric
processes than the emission of iodocarbons. At the same time the experiments
presented here show that the emission of iodocarbons is not linked to the
formation of I2 at the air–water interface (Martino et al., 2009),
since no correlation between I2 emissions or O3 mixing ratio and
iodocarbon emissions was observed.
Calculated emissions for the background, P. glacials and M. helysia were 8.32 × 105, 1.47 × 106
and 2.40 × 106 molecules cm-2 s-1, respectively.
Modelled emissions calculated using the kinetic model of the aqueous
interfacial layer by Carpenter et al. (2013) for the iodide concentration
measured were 1.16 × 106, 1.67 × 106 and 2.91 × 106 molecules cm-2 s-1, respectively. The measured and modelled values agree
well, showing that the model is able to predict emissions for natural iodide
concentrations.
Function of the change in the total I2 emissions in relation
to the amount of consumed ozone.
Figure 4 shows the change in the I2 emission
rate plotted versus the consumed ozone for the four different samples. This
was done to see whether ozone depletion in the flow chamber is mainly driven
by the iodide or whether other factors are important. The graph shows that the
ozone depletion correlates with the enhancement in the I2 emission
rate for the two diatom samples and for the background. Therefore, the ratio
of the formation of I2 to the amount of O3 consumed – calculated as
R(I2) = n(I2)/n(O3), with n(I2) = amount of
I2 formed and n(O3) = amount of O3 consumed during the
experiment – was used to determine the dependence of I2 formation on
O3. R(I2) has a maximum value of 1, which, referring to Eqs. (1–5),
indicates that every molecule of ozone which is consumed produces one
molecule of I2. The formation ratio for the background sample was the
highest with R(I2) = 0.14 ‰, followed by the
samples of M. helysia with R(I2) = 0.08 ‰ and P. glacialis
R(I2) = 0.07 ‰. This means that a higher degree of biological
activity of the sample decreases the formation ratio. The decrease of
I2 emission in the surface reaction of ozone with iodide was also seen
by Carpenter et al. when turning from iodide solutions to seawater, which
contains more organic substances (Carpenter et al., 2013).
The plankton concentrate also depletes ozone, although there is no
enhancement in I2 emission. Therefore, another mechanism in ozone
depletion must be taking place, possibly induced by other ozone-reactive
substances formed or excreted from Phaeocystis sp. Another explanation is a reduced
release of I2 and a higher release of HOI, which was not measured in
this study. In fact, Carpenter and co-workers observed HOI to be the main iodine
compound released in their experiments, followed by I2 (Carpenter et al., 2013).
Conclusions
Different phytoplankton suspensions were treated with high and low ozone
levels. Halocarbons including bromoform, iodomethane, iodochloromethane and
diiodomethane were released from the suspensions independent of the ozone
level. The use of an evaporation standard in the aqueous phase indicated
that the emission rates of all gaseous organics were quite stable. The
iodide and iodate concentration in the liquid phase also showed no
dependence on the ozone level in the gas phase and were comparable to
concentrations in surface water in the open ocean. The emission flux of the
iodocarbons was lower compared to the calculated flux at a coastal,
kelp-rich site in Roscoff, France, an observation which emphasizes the
higher emission of iodocarbons from macroalgae compared to microalgae. The
emission rates of iodocarbons were also lower than the emission of I2,
confirming that I2 emissions from the remote ocean dominate over
organic iodine sources for the MBL (Jones et al., 2010; Lawler 2012;
Carpenter et al., 2013). The emission of I2 showed a dependency on the
ozone level in the air as well as on the iodide concentration in the sample
suspension, as has been found previously (e.g. Carpenter et al., 2013).
For the two diatom samples, M. helysia and P. glacialis, and the background sample, a
correlation was found for the I2 emission and the ozone consumption
during the experiment. The I2 emissions from the plankton concentrate,
taken in the Wadden Sea of Sylt, were lower than the other samples and
showed no dependence on the ozone levels. An explanation could be the lower
iodide concentration in the plankton sample, since iodide is the limiting
factor for the oxidative reaction. Another explanation may be the preferred
formation and emission of HOI when organic compounds are present in the
liquid phase. The experiments showed that different algae suspensions (M. helysia and
P. glaciales) are capable of emitting I2 by the reaction of ozone with dissolved
iodide at the air–water interface under natural conditions. However, it
remains unclear whether iodine emissions from aquatic systems can be fully
understood without the simultaneous measurement of HOI.