Surface water distribution
CHBr3 and CH2Br2
The equatorial Atlantic is a source of CHBr3 and CH2Br2 to
the atmosphere during the ACT season, and the correlations of their water
concentrations to biogenic parameters indicate biological formation.
CHBr3 and CH2Br2 correlated significantly, but weakly with
T Chl a, which is not an unusual feature (Abrahamsson et al., 2004a; Carpenter
et al., 2009; Liu et al., 2011; Hepach et al., 2014). It has been suggested
that CHBr3 is not produced directly from phytoplankton, but rather from
dissolved organic matter (DOM) present in sea water (Lin and Manley, 2012).
This was more closely investigated in laboratory experiments by Liu et al. (2015), who suggested that the weak in situ correlations of bromocarbons
with Chl a are a result of this indirect production pathway. The correlation
with certain phytoplankton groups may then be caused by the production of
phytoplankton-specific DOM. The very negative correlations of
bromocarbons with SST and positive correlations with SSS indicate a
relationship of bromocarbon abundance with processes within the cold and
nutrient-rich upwelled water of the EUC (Sect. 3.2), supported by the
T-S diagrams (Fig. 4). Weak, but significant negative correlations with
latitude (rs=-0.38 for CHBr3 and rs=-0.18 for
CH2Br2) and maximum values of the bromocarbons between 2 and
3∘ S, where EUC water reaches the surface, underline this
hypothesis. Although the correlation analysis of halocarbons with
phytoplankton groups cannot directly resolve production and loss processes
by algal activity, it is still an indicator for possible involvement of
these species in halocarbon production. Bromocarbon production might exceed
loss processes, which leads to the observed statistical link of CHBr3
and CH2Br2 to chrysophytes. Chrysophytes are to our knowledge not yet among observed
halocarbon producers in incubation and field studies. The strong negative
correlations of Prochlorococcus HL with CHBr3 and CH2Br2 have been observed
previously (Hepach et al., 2014). These significant negative correlations
can be explained by the large abundance of Prochlorococcus in warm water while bromocarbons
on the other hand are more correlated with the cooler water of the EUC,
which is richer in nutrients and chrysophytes, haptophytes and dinoflagellates.
CH3I and CH2I2
CH3I concentrations and wind speed were weakly anticorrelated during
MSM18/3. Richter (2004) interprets this as depletion of the surface
concentrations, when air-sea fluxes exceed the production rate during high
wind speed. There are two production mechanisms suggested for CH3I.
Previous studies (Richter and Wallace, 2004; Jones et al., 2010) have
attributed CH3I in the tropical ocean mainly to photochemical formation
based on the observations of Moore and Zafiriou (1994). In contrast to these
studies, indications for biological formation of CH3I were found in the
ACT region during our study. CH3I showed a weak negative correlation
with SST, significant correlations with the biologically produced CHBr3
and CH2Br2 (Table 2) and with T Chl a as biomass indicator, and no
correlation to global radiation. These imply a relationship with the
biologically active upwelled water. Elevated concentrations of CH3I
were found between 10 and 5∘ W during midday (see
CH3I in comparison to global radiation in Fig. 2), which could be a
result of photochemical formation. Thus we suggest that photochemistry and
biological production likely both played a role during MSM18/3.
Haptophytes correlated most significantly of the phytoplankton groups with CH3I
and have already been shown to produce CH3I both in the laboratory
(Itoh et al., 1997; Manley and de la Cuesta, 1997; Scarratt and Moore, 1998;
Smythe-Wright et al., 2010) and in the field (Abrahamsson et al., 2004b).
Correlations during MSM18/3 additionally indicate a possible involvement of
dinoflagellates and chrysophytes in the production of methyl iodide (Table 2). The importance of
oceanic CH3I production by Prochlorochcoccus is a matter of dispute. Brownell et al. (2010) report it to be a minor source,
in contrast to both Smythe-Wright et al. (2006) and Hughes et al. (2010, 2011). No evidence of involvement of
Prochlorococcus HL was found during MSM18/3.
The very low sea surface concentrations of CH2I2 with lowest
concentrations during the day can be explained by its fast photolysis (few
minutes lifetime in surface sea water) (Jones and Carpenter, 2005; Martino
et al., 2005). Although CH2I2 is generally assumed to be of
biogenic origin in the open ocean (Moore and Tokarczyk, 1993; Yamamoto et
al., 2001; Orlikowska and Schulz-Bull, 2009; Hopkins et al., 2013), great
uncertainties remain as to which species are involved in its production.
During MSM18/3, indications were found for different source species than of
the other three compounds (chlorophytes and Prochlorococcus HL).
Water column distribution
Halocarbon maxima in the T Chl a maximum, attributed to their biological
production, are often observed from polar to tropical regions (Moore and
Tokarczyk, 1993; Moore and Groszko, 1999; Yamamoto et al., 2001; Quack et
al., 2004; Carpenter et al., 2007; Hughes et al., 2009). In contrast,
photochemical formation of CH3I can lead to surface maxima (Happell and
Wallace, 1996). During MSM18/3, maxima of halocarbons were not always found
in the T Chl a maximum. This does not contradict their biological production,
as the location of the T Chl a maximum is not necessarily the location of
highest biomass or primary production, but rather reflects the photoadaption
capability of the predominant phytoplankton groups (Claustre and Marty,
1995). Unfortunately, neither biomass nor primary production was measured
during the cruise. Additionally, halocarbons could be produced by
phytoplankton groups that are not in the maximum of the biomass distribution
in the water column, and the location of the halocarbon maximum might be
more determined from their sink processes than from their production.
Surprisingly, the time of day, influencing sink and production processes,
seemed to play a minor role for the shape of the profiles for all four
compounds (see the location of the CTD stations in Fig. 2).
CHBr3 and CH2Br2
In contrast to their similar occurrence in the surface, CHBr3 and
CH2Br2 showed different distributions in the water column (Fig. 5). Strong indications for biological sources of CHBr3 exist in the
PCA, and chrysophytes as potential source group are in agreement to the surface water
observations (Table 2, Fig. 5). Maximum CH2Br2 concentrations
were occasionally found below the CHBr3 maxima, which have already been
observed in the Mauritanian upwelling (Quack et al., 2007b). The deeper
maxima may be either due to an additional source of CH2Br2 such as
the biologically mediated conversion of CHBr3 (Hughes et al., 2013) or
to a faster degradation of CHBr3 than of CH2Br2 at depth.
Sinks for CHBr3 and CH2Br2 in tropical surface waters include
very slow hydrolysis (hundreds to thousands of years; Mabey and Mill, 1978)
and slow halogen substitution (5 years; Geen, 1992). Photolysis, which has
been suggested to be faster for CHBr3 (9 years with a mixed layer of
100 m for CHBr3) than for CH2Br2 (Carpenter et al.,
2009) would be of more significance in the surface layer. A faster
degradation of CHBr3 in greater depths is also somewhat contrary to the
observed very fast bacterial degradation of CH2Br2 with a
half-live of 2 days (Goodwin et al., 1998). An additional source for
CH2Br2 that involves CHBr3 therefore seems more plausible. At
four of the 13 stations, indications for the additional source were found.
There, maximum CH2Br2 concentrations were found below CHBr3,
which could be the result of its faster conversion to CH2Br2 than
its production. CH2Br2 in denser water is also co-located with
Prochlorococcus LL, which might be involved in the CHBr3-conversion.
CH3I and CH2I2
CH3I was usually elevated in the top 30 m of the water column apart
from three profiles, where maximum concentrations were found between 30 and
60 m. The surface maxima, as seen in the T-S diagram (Fig. 4), support the
photochemical formation of CH3I (Happell and Wallace, 1996). Deeper
maxima could also arise if the sea-to-air flux exceeds the photochemical
production. However, the low wind speed during the cruise (Sect. 3), the
relationship with biological parameters, and the partly co-located maxima
with the other three biogenic halocarbons (Figs. 3, 5) also point to
a direct production of CH3I from phytoplankton. These include
dinoflagellates as indicated by the correlations and the PCA (Fig. 5).
CH2I2 was always depleted in the surface with respect to the
underlying water column as a result of its strong photolysis (Jones and
Carpenter, 2005; Martino et al., 2006). It was frequently elevated below the
T Chl a maximum and below the base of the mixed layer (Fig. 3) in contrast
to previous studies (Moore and Tokarczyk, 1993; Yamamoto et al., 2001). The
similarity in its distribution to CH2Br2 (Figs. 4, 5)
could indicate similar production and sink processes at depth. Bacterial
formation of CH2I2 (Fuse et al., 2003; Amachi et al., 2005) in the
upper thermocline could also be an additional source for this compound.
Alternatively, CH2I2 may not degrade as quickly as CHBr3 and
CH3I in greater depths, which would lead to its accumulation below the
mixed layer.
Factors contributing to halocarbon emissions from the mixed layer
Halocarbon emissions into the atmosphere depend strongly on the mixed layer
budget of these compounds, which is determined by their sources and sinks.
It is unclear where the main halocarbon production occurs. It has been
suggested that it mainly takes place in the subsurface T Chl a maximum (Quack
et al., 2004; Martino et al., 2006), whereas other model studies assume
production of, e.g., CHBr3 to be coupled to primary production in the
whole water column (Hense and Quack, 2009). Assuming production of
halocarbons takes place mainly in the T Chl a maximum, which is often located
below the mixed layer, diapycnal fluxes from below the thermocline will be
the most important source for mixed layer halocarbons.
Transport and loss processes in the mixed layer
To evaluate the significance of halocarbon production below the mixed layer
for emissions into the atmosphere, production, loss and transport processes
have to be considered. The diapycnal fluxes of the four halocarbons were
calculated from 13 halocarbon profiles and parallel measurements of eddy
diffusivity (Sect. 4.3). The data are characterized by a low depth
resolution of the halocarbons within the water column and a short validity
of the diffusion coefficients, which make the diapycnal fluxes subject to
some uncertainties. Given that the depth profiles measured during MSM18/3
agree well to previous studies from the tropical ocean (Yamamoto et al.,
2001; Quack et al., 2004), a general idea of the significance of diapycnal
fluxes for the mixed layer budget of halocarbons can be obtained. The
chemical loss rates are estimated from published data which include
hydrolysis, halogen substitution and photolysis. The half-lives of
CHBr3 and CH2Br2 due to hydrolysis are hundreds to thousands
of years (Mabey and Mill, 1978), while for CH3I, the half-life due to
hydrolysis ranges from 1600 days at 25 ∘C to 4000 days at
5 ∘C (Elliott and Rowland, 1995). The half-life of CHBr3
with respect to photolysis is 9 years assuming a mixed-layer depth of 100 m
and is potentially longer for CH2Br2.
Liu et al. (2011) calculated the half-life of CHBr3 due to photolysis
in a coastal mixed layer of 5 m to be only 82 days. Mixed layers during
MSM18/3 were from surface down to 49 m, photolysis of bromocarbons in the mixed
layer will lead to half-lives of several months. Sea-to-air flux is the most
significant sink for CHBr3 and CH2Br2 from the mixed layer.
Mean half-lives of 8 days were calculated for both compounds during MSM18/3,
based on the fluxes (Sect. 4.3.1) and the mixed layer depths during the
cruise (Table 3). We consider a very short timescale of 1 h for our budget
calculations due to the validity of the diapycnal flux coefficients, while
the general findings of our calculations are also valid for a longer timescale. As the sink from the mixed layer due to sea-to-air fluxes is a
magnitude larger than the other mentioned sinks, we will neglect them in our
estimates for CHBr3 and CH2Br2 as they do not play a large
role. Photolysis of CH3I is very slow in comparison to halide
substitution (Zika et al., 1984). The latter is suggested to be an important
sink in the tropical ocean during low wind speeds (Jones and Carpenter,
2007), while large wind speeds favour sea-to-air fluxes as main sink (mean
half-life of 8 days during MSM18/3). All three sink processes are included
in our budget estimates using the rates published by Elliott and Rowland (1993). For CH2I2, photolysis is the most significant sink in
surface water (Jones and Carpenter, 2005). In our calculations, losses of
CH2I2 due to photolysis were calculated according to Martino et al. (2006) with a photon flux calculated from the NASA COART model (Jin et
al., 2006), a T Chl a concentration of 0.4 µg L-1, absolute
quantum yields from Martino et al. (2006), and absorption cross sections
determined by Jones and Carpenter (2005).
Mixed layer budget of halocarbons during MSM18/3
In the following section, the results of the halocarbon budget calculations
are presented. The total mixed layer concentrations were calculated at every
station considering a water column with a volume of 1 × 1 × zML
m3. Assuming that halocarbons are only produced below the
mixed layer, the following relationship (Eq. 4) is valid for the steady
state concentration Chal, with Fdia and Fadv as the source terms
from diapycnal fluxes and advection, while Sas (Fig. 6) and Sch
represent the loss terms sea-to-air flux and chemical sinks as described in
the previous section:
Chal=Fdia+Fadv-Sas-Sch.
Sas is the main sink term for CHBr3, CH2Br2 and
CH3I during MSM18/3 (Table 6). On the short timescales considered
here, diapycnal fluxes of CH3I, which can reduce the mixed layer by
around 5 pmol per hour (Table 5), compete with the loss due to chloride
substitution (Sch). For CH2I2, Sch (photolysis) is about
10 times higher than Sas, and reduces the mixed layer budget by 24 %
after 1 h. In total, diapycnal fluxes (Fdia) into the mixed layer were
not sufficient to account for the losses of all four compounds from the
mixed layer (Table 6). The discrepancies with respect to the total mixed
layer are 169 (CH2Br2), 255 (CH3I), 269 (CHBr3) to 8382
(CH2I2) pmol h-1, which are small compared to the total
amount of halocarbons in the mixed layer (CHBr3 – 0.17 %,
CH2Br2–0.19 %, CH3I–0.34 %, CH2I2–13.11 %). Possible reasons for the observed discrepancies are evaluated
in the following. Advection of the missing halocarbons, Fadv, likely
does not play a large role for CH2Br2, CH3I and
CH2I2, since mean mixed layer concentrations of these compounds
were rather homogeneous in the whole region. Thus, only for CHBr3, with
more variable concentrations, advection may transport significant amounts
from one location to another. In addition, halocarbon maxima were found
within the mixed layer, which may either result from a mixed layer that is
not well mixed or halocarbon production is faster than mixing in the mixed
layer. According to the temperature and salinity profiles during the whole
cruise (Fig. 3), the mixed layer was very well mixed. Consequently,
production in the mixed layer is the most likely process balancing the
missing halocarbons (Table 6) as diapycnal fluxes and advection play minor
roles. The maxima that occasionally evolve in the mixed layer suggest that
production of halocarbons is rapid, but may vary with depth. The mixed layer
production term, here called PML, has to be included in the budget
calculation of Eq. (4):
Chal=Fdia+Fadv-Sas-Sch+PML.
The relative production of halocarbons in the mixed layer is likely largest
for CH2I2, because its largest discrepancy arises from its rapid
photolysis (up to 24 % loss in 1 h; Table 6). This is in agreement to
earlier studies investigating macroalgal production, proposing larger
release rates of CH2I2 than of CHBr3, CH2Br2 and
CH3I (Klick and Abrahamsson, 1992; Carpenter et al., 2000).
Total mixed layer budget of each halocarbon, potential sinks and
sources (box size 1 × 1 × zML m3). The upper
four rows indicate cases where diapycnal fluxes act as sources, while the
lower four rows summarize the budget for the cases where the diapycnal
fluxes were sinks for the mixed layer budget. “Other sinks” is halogen
substitution for CH3I and photolysis in case of CH2I2. The
negative numbers indicate sinks for the budget.
Compound
zML
Total ML
Air-sea fluxes
Diapycnal fluxes
Other sinks
Total after
Difference
Unit
[m]
budget [pmol]
(Sas) [pmol h-1]
(Fdia) [pmol h-1]
(Sch) [pmol h-1]
1 h [pmol]
[pmol]
CHBr3
24
157543
-274
5
157274
-269
Diapycnal fluxes
CH2Br2
29
90058
-172
3
89889
-169
as source
CH3I
26
75263
-257
2
0
75004
-255
CH2I2
28
63947
-78
13
-8317
55565
-8382
CHBr3
36
417098
-1186
-30
415882
-1216
Diapycnal fluxes
CH2Br2
27
99604
-236
-2
99366
-238
as sink
CH3I
29
137560
-420
-5
0
137135
-425
CH2I2
29
106587
-35
-2
-4977
101573
-5014
Production rates of halocarbons
From the budget calculations, described in the previous section, potential
production rates PML for the mixed layer are determined for each
station. The mean production rates show large standard deviations (Table 7),
including the variability and uncertainties in the estimated production
rates. Production rates are 34 ± 65 (CHBr3), 10 ± 12
(CH2Br2), 21 ± 24 (CH3I), and 384 ± 318 pmol m-3 h-1 (CH2I2). These are the first estimated
production rates of CHBr3 and CH2Br2 for tropical
phytoplankton species. For comparison to other studies, the production rates
from this study are converted to rates per µg T Chl a (reported in
Tables 3 and 4), which results in mean (±standard deviation)
production rates of 2.5 × 10-3 ± 4.5 × 10-3 (CHBr3),
8.4 × 10-4 ± 1.0 × 10-3 (CH2Br2), 2.2 × 10-3 ± 3.0 × 10-3 (CH3I) and
3.3 × 10-2 ± 3.3 × 10-2 pmol [µg T Chl a]-1 h-1 (CH2I2).
Theoretical mean production rate of the four halocarbons in the
equatorial mixed layer with the standard deviation.
Compound
Production rate
Standard deviation
Production rate per T Chl a
[pmol m-3 h-1]
[pmol m-3 h-1]
[pmol [µg T Chl a]-1 h-1]
CHBr3
34
65
2.5 × 10-3
CH2Br2
10
12
8.5 × 10-4
CH3I
21
24
2.2 × 10-3
CH2I2
384
318
3.3 × 10-2
Comparison to previously reported rates – CHBr3 and
CH2Br2
Tokarczyk and Moore (1994) and Hughes et al. (2013) determined production
rates from polar algae in laboratory studies ranging between 2 × 10-3
and 2.1 × 10-2 pmol [µg Chl a]-1 h-1 on average for
CHBr3, depending on the growth phase, which is in the range of our
calculated rates. Production rates for CH2Br2 of on average 2.1–4.2 × 10-3 pmol [µg Chl a]-1 h-1 were much higher than
the ones calculated in our study (Tokarczyk and Moore, 1994). Karlsson et al. (2008) published production rates of 2.6–9.3 × 10-2 pmol
[µg Chl a]-1 h-1 for CHBr3 (depending on the time
of day) and 5 × 10-4–3.6 × 10-3 pmol [µg Chl a]-1
h-1 for CH2Br2 from an in situ study in the Baltic Sea during
a cyanobacterial bloom. Liu et al. (2011) calculated 417 (CHBr3) and
258 pmol m-3 h-1 (CH2Br2) for the subtropical and
temperate eastern US coast, which are tenfold higher than the production
rates determined from our study (Table 7). The differences between these
studies and ours may have several origins. Taking an average production rate
for the total mixed layer during MSM18/3 does not take a potential variable
production with depth into account. Second, the different production rates
determined in the monocultural studies (Tokarczyk and Moore, 1994; Hughes et
al., 2013) show large variations between different types of microalgae.
Third, the indirect estimates during MSM18/3 are afflicted by the
uncertainties in the individual budget terms, which are also expressed in
the large standard deviations.
Comparison to previously reported rates – CH3I and
CH2I2
Production rates of CH3I determined from Prochlorococcus vary significantly from 5.8 × 10-4 to 9.4 × 10-2 pmol [µg
Chl a]-1 h-1 (Smythe-Wright et al., 2006; Brownell et al., 2010). Hughes et al. (2011) suggested this variability to be caused by different cell states,
e.g. healthier cells producing less CH3I. While Scarratt and Moore
(1999) determined rates from 8.3 × 10-3–5.0 × 10-2 pmol
[µg Chl a]-1 h-1 from a red microalgal species, Karlsson
et al. (2008) reported a rate of 1.0 × 10-2 pmol CH3I [µg
Chl a]-1 h-1 from a cyanobacterial bloom in the Baltic Sea,
which is at the higher end of the range mentioned here. Our estimates lie
well within these cited ranges of phytoplankton production rates
and are thus a reasonable assumption for the CH3I production strength
of tropical algae (see Sect. 5.1.2).
In contrast to the other three halocarbons, very few studies have actually
determined production rates of CH2I2 from phytoplankton.
CH2I2 was shown to be produced in comparatively larger
concentrations than other halocarbons, but generally from fewer species. Six
polar and temperate diatom species were tested, of which only two produced
CH2I2 (Moore et al., 1996). Martino et al. (2006) assumed a
theoretical production rate of 17 000 pmol m-3 h-1 in the tropical
equatorial Atlantic. These were calculated from previously reported
CH2ClI fluxes based on the assumption that CH2ClI is mainly formed
during the photolysis of CH2I2 and that CH2I2 is only
produced in the T Chl a maximum. This rate appears very large in comparison to
our estimate and in comparison to the production rates of the other
halocarbons. We showed evidence that CH2I2 is not only produced
within the T Chl a maximum but in the whole mixed layer, thus, lower average
production rates seem more plausible. CH2I2 together with
CH2ClI have been suggested to be equally important carriers of
organoiodine into the troposphere (Saiz-Lopez et al., 2012), hence it is
important to determine specific phytoplankton production rates of
CH2I2 in future studies.
Our calculated production rates of CHBr3, CH2Br2 and
CH3I lie well within the ranges of several laboratory and field studies
of mostly temperate and polar algae, suggesting production from tropical
algae to be similarly significant. CH2I2 was shown to be produced
in larger rates than the other three compounds, but very rapid photolysis
leads to lower sea surface concentrations of this compound. However,
considering the large ranges in reported production rates of CHBr3,
CH2Br2, CH3I and the lack of studies concentrating on
CH2I2, more incubation experiments are severely needed to
constrain in situ production rates of tropical algae. This information is
crucial to evaluate the significance and contribution of the tropical ocean
with respect to halogen transport into the troposphere, and finally into the
stratosphere. Understanding the fate of halocarbons within the water column
is an important task to estimate their distribution and emissions from the
future ocean.