Introduction
In the mid-1990s, the recognition that the known CH3Cl sources,
mainly biomass burning and marine emissions, are insufficient to balance the
known atmospheric sinks (Butler, 2000). This motivated intense research on
potential terrestrial sources. Today, it is common thinking that large
emissions from tropical rainforests (Monzka et al., 2010; Xiao et al., 2010;
Carpenter et al., 2014) can close this gap. Several model studies revealed a
strong tropical CH3Cl source in the range of 2000 Gg a-1
(Xiao et al., 2010; Yoshida et al., 2006; Lee Taylor et al., 2001).
Particular support for a strong tropical rainforest source arose from
observations of elevated CH3Cl concentrations in the vicinity of
tropical rainforests (Yokouchi et al., 2000), greenhouse experiments
(Yokouchi et al., 2002), several field measurements in tropical rainforests
(Saito et al., 2008, 2013; Gebhardt et al., 2008; Blei et al., 2010) and from
carbon stable isotope mass balances (Keppler et al., 2005; Saito and
Yokouchi, 2008). The majority of CH3Cl in tropical rainforests
((2000±600) Gg a-1) is thought to originate from higher plants
(Monzka et al., 2010; Xiao et al., 2010; Yokouchi et al., 2000; Saito and
Yokouchi, 2008). A minor fraction of about 150 Gg a-1 may be emitted
from wood rotting fungi (Monzka et al., 2010; Xiao et al., 2010; Carpenter et
al., 2014). Further emissions from senescent leaf litter (Keppler et al.,
2005) may substantially contribute to this source, but this has not yet been
confirmed in field studies (Blei et al., 2010). On a global scale, biomass
burning (400 to 1100 Gg a-1) and surface ocean net emissions (140 to
640 Gg a-1) are further important sources (Monzka et al., 2010; Xiao
et al., 2010; Carpenter et al., 2014). Chloromethane from higher plants has
an average stable isotope signature (13C/12C ratio,
δ13C value) of (-83±15) ‰ (Saito et al., 2008;
Saito and Yokouchi, 2008). Compared to the other known sources with
δ13C values in the range -36 ‰ to -62 ‰
(Keppler et al., 2005; Saito and Yokouchi, 2008), the tropical rainforest
source is exceptionally depleted in 13C, making stable isotope
approaches particularly useful for better constraining CH3Cl flux
estimates.
The isotopic composition of tropospheric CH3Cl links the isotopic
source signatures to the kinetic isotope effects (KIEs) of the sinks. The
primary CH3Cl sink is its oxidation in the troposphere by OH and Cl,
accounting for about 80 % of total losses (Monzka et al., 2010; Xiao et
al., 2010; Carpenter et al., 2014). Further sinks comprise soil uptake and
loss to the stratosphere (Monzka et al., 2010; Xiao et al., 2010; Carpenter
et al., 2014). An accurate determination of the KIEs of the main
tropospheric sink reactions (CH3Cl+OH, CH3Cl+Cl) is
crucial for constraining the tropical rainforest source from an isotopic
perspective. A previous study (Gola et al., 2005) revealed large
ε values of (-59±8) ‰ and (-70±10) ‰ for the reaction of CH3Cl with OH and
Cl, respectively, which supported the hypothesis of large emissions from
tropical rainforests (Keppler et al., 2005; Saito and Yokouchi, 2008). In
particular, the ε value for the reaction with OH is much larger
in comparison to previously reported ε values for the reaction
of OH with methane (Saueressig et al., 2001) and other hydrocarbons (Rudolph
et al., 2000; Anderson et al., 2004). We thus performed photochemical
degradation experiments of CH3Cl in a 3500 dm3 Teflon smog chamber
using established radical generation schemes (see method section for
details) to reassess the KIEs for the reaction of CH3Cl with OH and Cl.
For validation purposes, we further determined the known KIEs for the same
reactions of methane (CH4).
In the next step, we used the seasonal variations in the mixing ratios
(Prinn et al., 2000) and isotopic composition (Thompson et al., 2002;
Redeker et al., 2007) of tropospheric CH3Cl to further assess the
reliability of the obtained KIEs. This was done with a simple two-box model,
dividing the atmosphere into a Northern and a Southern Hemisphere and using
a simplified emission scheme. The same model was then used to constrain the
tropical rainforest source from an isotopic perspective. We finally improved
previous bottom-up estimates of the tropical rainforest source using carbon
density maps of the tropical rainforest instead of coverage area.
The kinetic isotope effect (ε) for the reaction of
CH3Cl with OH and Cl
Materials and methods
Smog chamber
The smog chamber set-up and the experimental conditions are the same as
recently described in Keppler et al. (2018). The samples for the carbon
isotope analysis were taken from the same experiments described therein.
Briefly, the isotope fractionation experiments were performed in a (3500±100) dm3 Teflon smog chamber. The chamber was continuously flushed
with purified, hydrocarbon-free zero air (zero-air generator, cmc
instruments, <1 nmol mol-1 of O3,
<500 pmol mol-1 NOx, <100 nmol mol-1 of
CH4) at a rate of 0.6–4 dm3 min-1 to maintain a slight
overpressure of 0.5–1 Pa logged with a differential pressure sensor
(Kalinsky Elektronik DS1). A Teflon fan inside the chamber ensured constant
mixing throughout the experiments. NO and NOx were monitored
on a routine basis with an Eco Physics, CLD 88p chemiluminiscence analyser
coupled with an Eco Physics photolytic converter, PLC 860. Ozone was
monitored by a chemiluminescence analyser (UPK 8001). Initial CH3Cl
mixing ratios were between 5 and 14 µmol mol-1.
Perfluorohexane (PFH) was used as an internal standard with initial mixing
ratios of (25±3) µmol mol-1 to correct the resulting
concentrations for dilution. The mixing ratios of CH3Cl and PFH
were monitored by GC–MS (Agilent Technologies, Palo Alto, CA) with a time
resolution of 15 minutes throughout the experiments. The stability of the
instrument was regularly checked using a gaseous standard (5 mL of
100 µmol mol-1 CH3Cl in N2). Mixing
ratios of methane and CO2, used as an internal standard in the
methane degradation experiments, were measured with a Picarro G221i cavity
ring-down spectrometer. Prior to the experiments, the instrument was
calibrated with pressurized ambient air from a tank obtained from the
Max-Planck-Institute for Biogeochemistry in Jena, Germany (CO2
mixing ratio of (394.6±0.5) µmol mol-1, methane mixing
of (1.752±0.002) µmol mol-1). OH radicals were generated
via the photolysis of ozone (about 2000 nmol mol-1 for CH3Cl
and about 10000 nmol mol-1 for CH4) at 253.7 nm in the
presence of water vapour (Relative humidity = 70 %). This is a
well-established efficient method for OH radical generation (Cantrell et al.,
1990; DeMore, 1993). In the CH3Cl+OH experiments, initially
2000 µmol mol-1 of H2 was added for scavenging
chlorine radicals originating from the photolysis or oxidation of formyl
chloride (HCOCl) occurring as an intermediary in the reaction cascade (Gola
et al., 2005). To obtain an efficient OH formation, Philips TUV lamps (1×55 W for CH3Cl, 4×55 W for CH4) were
welded in Teflon film and mounted inside or around the smog chamber. Atomic
chlorine (Cl) was generated via photolysis of molecular chlorine
(Cl2) at a relative humidity of less than 1 % by a solar
simulator (Behnke et al., 1988) with an actinic flux comparable to the sun at
midsummer in Germany and a photolysis frequency of J=1.55×10-3 s-1 for Cl2 (Buxmann et al., 2012). A more detailed
description of the smog chamber set-up is provided in the Supplement and has
recently been published elsewhere (Wittmer et al., 2015; Keppler et al.,
2018).
Sampling and carbon isotope determination
From each experiment 10 to 15 canister samples (2 dm3 stainless steel,
evacuated <1.3×10-3 Pa and baked out at 250 ∘C for
2 h) and 10 to 15 adsorption tube samples were taken at regular time
intervals for subsequent analysis of carbon isotope ratios. The adsorption
tubes were made of stainless steel (1/4 inch outer diameter, 7 inch length)
and filled with 77 mg Carboxen 1016®,
215 mg Carbopack X 569®, 80 mg
Carboxen® 1003 and 9 mg
Tenax® TA in order of the sampling flow
direction. The adsorption tube samples and one set of canister samples from
the CH3Cl degradation experiments were analysed by 2D-GC–IRMS/MS
at the University of Hamburg using the method of Bahlmann et al. (2011). This
method has been shown to be free of interference from other compounds. The
precision and reproducibility of the δ13C measurements based
on standards were ±0.6 ‰ (n=18) on the 1σ level. In
order to assure compliance with VPDB scale, a single-component standard of
CH3Cl (100 µmol mol-1 in nitrogen, Linde Germany)
was calibrated against a certified CO2 reference standard (Air
Liquide, Germany, (-26.8±0.2) ‰) and a solid standard (NIST
NBS 18, RM 8543) after offline combustion and analysis via a dual inlet (DI).
The results from the DI (n=6) were (-37.2±0.1) ‰ for
CH3Cl. The respective δ13C values from the
GC-GC–IRMS, measured against the machine working gas (Air Liquide, Germany,
(-26.8±0.2) ‰) were (-36.1±0.2) ‰, resulting in
an offset (DI – 2D-GC–IRMS) of -1.1 ‰ for CH3Cl.
The canister samples were analysed at the University of Heidelberg using a
cryogenic pre-concentration system coupled to a GC-C-IRMS
system, developed for δ2H measurements of CH3Cl
(Greule et al., 2013). A combustion reactor filled with copper (II) oxide at
850 ∘C was used to analyse δ13C. The precision and
reproducibility of these δ13C measurements based on a
CH3Cl working standard were ±0.47 ‰ (n=47) on the
1σ level. The sampled CH3Cl amounts varied between 0.8 and
15 nmol. Both methods were linear over the whole range of sampled
CH3Cl amounts. The δ13C values measured in both
laboratories generally agreed within ±1.3 ‰ on the 1σ
level. This range is somewhat larger than expected from error propagation and
may result from small additional errors of scale adding to the uncertainty.
For the purpose of this study no attempts were made to adjust the measured
δ13C values. CH4 carbon isotope ratios were only
analysed at the University of Heidelberg.
Calculation of ε
The carbon isotope ratios are reported in the δ notation relative to
the VPDB scale (Vienna Pee Dee Belemnite), and the kinetic isotope effect
(KIE, symbol ε) is reported in ‰. We
applied an orthogonal regression model (Danzer et al., 1995) to derive the
kinetic isotope effect for each experiment from the slope of the Rayleigh
plot:
ε×lnft=lnδ13Ct+1δ13C0+1,
with ε being the kinetic isotope effect, ft being the
residual CH3Cl fraction at time t, δ13C0 being the
initial carbon isotope ratio of the substrate ‰,
δ13Ct being the carbon isotope ratio of the substrate
‰ at time t. To account for the dilution from the airflow through the chamber, the residual fraction (ft) has been
calculated from the mixing ratios of CH3Cl and the inert tracer, PFH,
as follows:
ft=CH3Clt×PFH0CH3Cl0×PFHt.
Here, [CH3Cl] and [PFH] denote the respective concentrations, and the
indices t and 0 refer to time t and zero. The uncertainty for
ft ranged from 1.4 % to 1.8 % on the 1σ level.
Results of the CH3Cl degradation experiments
In total, we performed six degradation experiments and two control
experiments within this study. To perform the degradation experiments within
a day, the experimental conditions were modified as indicated in Table 1. For
the OH experiments in the presence of CH4, the light intensity was
increased from 55 to 220 W, and the steady-state ozone mixing ratios were
increased from about 620 nmol mol-1 to about 3570 nmol mol-1. Under
these experimental conditions, typically 70 % to 80 % of the initial
CH3Cl and CH4 were degraded within 6 to 10 h. A more detailed
discussion of the experimental conditions with respect to the OH yields and
degradation rates is provided in the Supplement. Further, the reader is
referred to Keppler et al. (2018), who reported on the hydrogen isotope effects
from these experiments.
Experimental conditions of the degradation experiments with OH. The
O3 mixing ratios are average steady-state mixing ratios throughout
the experiment, the Cl2 mixing ratios refer to the initial mixing
ratios at the beginning of each photolysis sequence.
Experiment
Reactant
Oxidant
O3, Cl2
Irradiation
H2
Relative humidity
T
OH
µmol mol-1
µmol mol-1
µmol mol-1
%
∘C
cm-3
1 and 2
CH3Cl
5, 10
OH
0.62
1×55 W, λmax=254 nm
2000
65
20.7
2.9×109
3
CH3Cl
0.13
OH
0.62
1×55 W, λmax=254 nm
2000
<1
20.6
8.7×107
4
CH3Cl, CH4
13, 5
OH
3.7
4×55 W, λmax=254 nm
2000
65
20.4
1.6×1010
5
CH4
5
0
4×55 W, λmax=254 nm
72
20.3
6
CH4
6
OH
3.7
4×55 W, λmax=254 nm
72 to 75
20.3
1.6×1010
7
CH3Cl
10
Cl
2 to 10
7×1200 W 300–700 nm
<1
20.7
8
CH4
5
Cl
2 to 10
7×1200 W 300–700 nm
<1
20.5
Prior to each degradation experiment, we monitored the ratio of CH3Cl
and perfluorohexane (PFH) for at least 2 h to assess potential side reactions
and unwanted losses of CH3Cl. For the experiment with chlorine, this
was done under dark conditions in the presence of 10 µmol mol-1
Cl2. For the OH experiments, this was either done in the absence of
light or ozone. None of these tests revealed an indication of a measurable
loss (1.4 % to 2.1 %) of CH3Cl and thus for any biasing side effects
or reactions. In the CH4 degradation experiments, CO2 was used as
an internal standard to correct the CH4 mixing ratios for dilution. A
control experiment over 9 h, carried out with a dilution flow of 4 dm3 min-1 of zero air, revealed a slope (-0.00118±0.00001) min-1 for CH4
loss and a slope of (-0.00117±0.000007) min-1 for the CO2 loss. This corresponds to a
dilution flow of (4.1±0.1) dm3 min-1, which is in good
agreement with the pre-set dilution flow (the major uncertainty in this
calculation is the exact volume of the chamber). During this control
experiment, the dilution-corrected mixing ratio of CH4 changed by less
than 0.2 %.
Change in δ13C over the extent of the reaction (a)
and corresponding Rayleigh plots (b) from the CH3Cl
degradation experiments. Filled symbols and regression lines refer to the
data from Heidelberg; open symbols and dashed regression lines show data from
Hamburg. The colours refer to different degradation experiments (black is
experiment 1, red is experiment 2, blue is experiment 4 and green is experiment 7).
Errors in ζ were ±2 % on the 1σ level. Errors in
δ13C, derived from the regression analysis, ranged from
±0.4 ‰ to ±1.4 ‰ on the 1σ level.
In our study, photolysis of ozone (620 nmol mol-1 steady-state mixing
ratio) in the absence of water vapour (relative humidity <1 %) but
with 2000 µmol mol-1 H2 (experiment 3) resulted in
a CH3Cl degradation of less than 3 % over 10 h and no
measurable change in the isotopic composition of CH3Cl
(-46.8 ‰ at the beginning and -46.1 ‰ after 10 h)
because of the insufficient OH yield. The reaction rate constants of
O(1D) with H2 and H2O at 298 K are 1.1×10-10 and 2.2×10-10 cm3 s-1, respectively
(Burkholder et al., 2015). At a relative humidity of 70 % (corresponding
to 16 000 µmol mol-1), the reaction with H2O is by
far the main pathway along which OH forms (with the H2 pathway
contributing less than 4 % to the OH yield). This is consistent with the
previous study, where ozone levels of 300 µmol mol-1 were
required for a sufficient OH production from H2 (Gola et al., 2005;
Sellevåg et al., 2006). For this experiment, the partial lifetime of
CH3Cl with respect to OH can be estimated to be about 330 h.
Potential side reactions with O(1D) were not explicitly
investigated in our study but because of the reduced OH yield, this
experiment allows potential losses of CH3Cl to be constrained due
to the reaction with O(1D). In experiment 4, where both
CH3Cl and CH4 were present, the ratio of the measured
rate constants for the reaction of CH3Cl and CH4 with OH
was 5.8. This ratio agrees well with that of the recommended rate constants
of 5.6 at 298 K (6.3×10-15 cm3 s-1 for CH4
and, 3.5×10-14 cm3 s-1 for CH3Cl at 298 K;
Burkholder et al., 2015).
Summary of the kinetic isotope effects (ε) for the
reaction of CH3Cl and CH4 with OH and Cl from this study.
We used an orthogonal regression to calculate ε and the
respective uncertainties on the 1σ level for each experiment. In
experiment 4 the ε for methane has not been determined.
Hamburg
Heidelberg
ε
R2
ε
R2
Experiment 1
CH3Cl+OH
-12.1±0.6
0.95
-11.7±0.4
0.99
Experiment 2
-12.1±0.3
0.99
-10.5±0.3
0.99
Experiment 4
-10.4±0.4
0.99
-10.6±0.6
0.99
Experiment 7
CH3Cl+Cl
-10.3±0.7
0.96
-10.4±0.4
0.98
Experiment 5
CH4+OH
-4.7±0.2
0.99
Experiment 8
CH4+Cl
-59.0±1.3
0.99
Compilation of kinetic isotope effects (ε) for the
reaction of CH3Cl, CH4 and alkanes with OH and Cl.
Reaction
ε
Method
Reference
CH3Cl+OH
-58±10
smog chamber; O3+H2+ hv (254 nm), FTIR
Gola et al. (2005)
-44
theoretical at 298 K
Feilberg et al. (2005)
-3.6
theoretical
Jalili and Akhavan (2006)
-5±3
derived from field data
Thompson et al. (2002)
-11.2±0.8
smog chamber; O3+H2O +hv (254 nm), GC–IRMS
This study
CH3Cl+Cl
-70±10
smog chamber; Cl2 +hv (370 nm), FTIR
Gola et al. (2005)
-35
theoretical at 298 K
Feilberg et al. (2005)
-10.4±0.5
smog chamber; Cl2 +hv, GC–IRMS
This study
CH4+OH
-3.9±0.4
photoreactor; H2O2 +hv, GC–IRMS
Saueressig et al. (2001)
-5.4±0.9
photoreactor, O3+H2O +hv (254 nm), GC–IRMS
Cantrell et al. (1990)
-4.7±0.2
smog chamber; O3+H2O + hv (254 nm), GC–IRMS
This study
CH4+Cl
-58±2
smog chamber; Cl2 +hv FTIR
Sellevåg et al. (2006)
-66±2
photoreactor; Cl2 +hv; TDLAS
Saueressig et al. (1995)
-62±0.1
smog chamber; Cl2 +hv; DI-IRMS
Tyler et al. (2000)
-59±1.3
smog chamber; Cl2 +hv GC–IRMS
This study
C2H6+OH
-7.5±0.5
reaction chamber; H2O2 +hv; GC–IRMS
Piansawan et al. (2017)
R-CH3+OH
-18.7±5.2
reaction chamber; R-NO2, NO+hv; GC–IRMS
Anderson et al. (2004)
R-CH3+Cl
-18.6±0.3
reaction chamber; Cl2 +hv; GC–IRMS
Anderson et al. (2007)
The change in stable carbon isotope δ values of CH3Cl
(δ13C(CH3Cl)) with the extent of the reaction and the
corresponding Rayleigh plots of the CH3Cl degradation experiments
are shown in Fig. 1. The respective ε values, derived from the
slope of the Rayleigh plot, are summarized in Table 2. For the reaction of
CH3Cl with OH, we determined an ε value of (-11.2±0.8) ‰ (n=3) and for the reaction with Cl we found an
ε value of (-10.2±0.5) ‰ (n=1). The results from
both laboratories generally agreed within ±1.5 ‰ (1σ)
and showed no systematic difference. Variations in the initial mixing ratios
(5 to 13 µmol mol-1) and isotopic composition ((-47.0±0.5) ‰ and (-40.3±0.5) ‰) of CH3Cl in the
OH experiments had no significant effect on the determination of the isotope
effects. Furthermore, the increase in the light intensity and ozone mixing
ratios in experiment 4 had no effect on the isotope effects.
The ε values for the reaction of CH4 with OH and Cl,
determined for validation purposes, agreed reasonably well with the
previously published KIEs (Saueressig et al., 1995, 2001; Tyler et al.,
2000; Feilberg et al., 2005). For the reaction of CH4 with OH, we found
an ε value of -4.7 ‰, which is at the upper
end in terms of absolute magnitude of previously reported fractionation
factors, and for the reaction with Cl we found an ε value of
-59 ‰, which is more at the lower end of previously
measured KIEs (Table 3). Prior to the CH4 degradation experiment with
OH, we performed a control experiment (experiment 5 in Table 1) that revealed no
CH4 loss over 10 h. With this, we can exclude any interference from
reactive chlorine during the CH4–OH experiment. The larger isotope
effect for the reaction with OH found here might result from the reaction of
CH4 with O(1D). Cantrell et al. (1990), who also used
UV-photolysis in the presence of water as an OH source, reported an even
higher ε value of (-5.4±0.9) ‰ and
estimated that the reaction of CH4 with O(1D) (showing an
ε value of -13 ‰; Saueressig et al., 2001)
may have contributed about 3 % to the overall degradation. Saueressig et
al. (2001) reported an ε value of -3.9 ‰
for the reaction of CH4 with OH. With respect to this value, a
contribution of 9 % from the reaction with O(1D) is required to
explain the difference in terms of O(1D) loss.
Discussion of the CH3Cl degradation experiments
Our newly determined isotope effects for the reaction of CH3Cl with OH
and Cl are 5 to 6 times smaller than the previously reported ε values of (-59±10) ‰ for the reaction with OH
and of (-70±10) ‰ for the reaction with Cl (Gola
et al., 2005). In this section, we first discuss potential sources of error
in our study with particular respect to the differences between our study
and the Gola study and then provide a more comprehensive comparison of our
data with previous data. Gola et al. (2005) used a 250 dm3 electro
polished stainless steel chamber for their degradation experiments. We used
a 3500 dm3 smog chamber, made of FEP foil, for the CH3Cl
degradation experiments. The large volume of our smog chamber may result in
incomplete mixing and thus in an underestimation of the KIE due to transport
limitation. The effect of mixing on the observed KIE can be approximated
from the timescales of mixing and reaction according to the following
equation (Morgan et al., 2004; Kaiser et al., 2006):
ϵobs≈12εi×1+11+Q.
Here, εi is the intrinsic fractionation factor,
ϵobs is the observed fractionation factor and Q is the
ratio of the mixing time and reaction timescale (1/k). The chemical
lifetime of CH3Cl under the experimental conditions was in the
order of 6 to 8 h and the turnover of air inside the chamber occurred on
timescales of a few minutes. With a reaction timescale of 300 minutes and a
mixing timescale of 10 min we obtain εobs≈0.99×εi, making incomplete mixing an unlikely
source of error. Incomplete mixing would also have affected the determination
of the respective KIEs for CH4. With those values being within
previously reported ε values, we can exclude incomplete mixing
as a potential source of error in our experiments. In the Gola et al. (2005)
study, the mixing ratios and isotope ratios were determined with long-path
FTIR. In our study, the mixing ratios were determined by GC–MS and the
isotope ratios were measured by GC–IRMS in two different laboratories. Both
labs used different analytical set-ups, different sampling methods and
different standards. However, the results from both labs generally agree
within ±1.5 ‰ on the 1σ level and show no systematic
difference. As outlined before, different initial δ13C(CH3Cl) as well as different initial CH3Cl mixing ratios
had no significant effect on the determination of the isotope effects. This
makes analytical artefacts in our δ13C determination unlikely.
The Cl radical generation scheme was quite similar among both studies. Gola
et al. (2005) used narrowband photolysis of Cl2, employing a
Philips TLD-08 fluorescent lamp (λmax∼370) nm, whereas we used
broadband photolysis (300 to 700 nm), making this an unlikely source for the
discrepancy in between the isotope effects for the reaction of
CH3Cl with Cl.
In our study, OH was generated via UV photolysis of ozone (steady-state
mixing ratios of 0.62 and 3.6 µmol mol-1) in the presence of
water vapour (RH of 70 %) and 2000 µmol mol-1 H2, whereas
in the Gola study OH was generated in the absence of water vapour from the
reaction of O(1D) with H2 (2000 µmol mol-1) after
UV-photolysis of ozone (300 µmol mol-1). Due to the much lower
ozone mixing ratios employed in our study, the OH generation in the absence
of water vapour was not sufficient in our study. Both OH generation schemes
are well established. However, Cantrell et al. (1990), who used
UV-photolysis in the presence of water as an OH source, estimated that the
reaction of CH4 with O(1D) may contribute about 3 % to the
overall degradation. The higher ozone levels and the less efficient
conversion of O(1D) to OH in the Gola et al. (2005) study suggest an
overall higher transient O(1D) concentration compared to our
experiments. Anyhow, interference from the reaction with O(1D) is
less likely for CH3Cl than for CH4. The reaction rate for CH4
with O(1D) (1.7×10-10 cm3 s-1; Burkholder
et al., 2015) is 2.7×104 times larger than the respective reaction rate
for OH (6.3×10-15 cm3 s-1). In the case of
CH3Cl, the ratio is only 7.4×103 (2.6×10-10 and 3.5×10-14 cm3 s-1; Burkholder et al., 2015). Assuming a contribution of 9 % from
the reaction with O1D in the CH4 experiment, the reaction with
O(1D) should contribute less than 2.3 % to the observed CH3Cl
loss. In the CH3Cl control experiment, all experimental parameters
besides the relative humidity and hence the OH yields were comparable to
CH3Cl degradation experiments with OH. The CH3Cl loss of less than
3 % over 10 h can most likely be attributed to a reaction with OH,
originating from the reaction of O(1D) with H2. The δ13C values of CH3Cl were -46.8 ‰ at the
beginning and -46.1 ‰ after 10 h, which are
indistinguishable within the measurement uncertainty. This experiment makes
any biasing side reactions unlikely. In any case we can limit the loss from
potential side reactions to less than 3 %. In addition, none of our tests
prior each degradation experiment revealed an indication of a measurable loss
of CH3Cl. With this, we can safely exclude any measurable effect from
potential side reactions on the determination of the KIEs in our study.
A comparison of our data with previously measured and calculated isotope
effects for the reaction of CH3Cl, CH4 and other VOCs
with OH and Cl is provided in Table 3. In a follow-up study to Gola et
al. (2005), Sellevåg et al. (2006) attributed these exceptionally large
fractionation factors to higher internal barriers of rotation of the OH
radical compared to the CH4+OH reaction. Using variational
transition state theory, the authors calculated ε values of
-47 ‰ and -37 ‰ for the reaction of CH3Cl
with OH and Cl, respectively. However, a simultaneous theoretical study
provided an ε value of only -3.6 ‰ for the reaction
of CH3Cl with OH (Jalili and Akhavan, 2006). For C–H bond
breakage, Streitwieser's semi-classical limit for isotope effects is
-21 ‰ (Elsner et al., 2005), and for reactions involving hydrogen
radical transfer, an ε value of -15 ‰ has been
reported (Merrigan et al., 1990). Both values support a lower fractionation
factor. For the reaction of ethane with OH, which can be approximately
regarded as a substituted methane, an ε value of (-7.5±0.5) ‰ has been reported (Piansawan et al., 2017). One can estimate
an upper limit for the reactive site by multiplying δ13C with
the number of carbon atoms in the molecule (Anderson et al., 2004). This
leads to an upper limit of (-15.0±0.7) ‰ for the ε value at the reactive centre. In line with this, Anderson et al. (2004)
reported a group kinetic isotope effect of (-18.7±5.2) ‰ for
the reaction of primary carbon atoms of alkanes with OH. The same group
(Anderson et al., 2007) found a group kinetic isotope effect of (-18.6±0.3) ‰ for the respective reaction with Cl. Our smaller kinetic
isotope effect for the reaction of CH3Cl with OH and Cl is much
closer to these group-specific kinetic isotope effects than the previously
reported ones from Gola et al. (2005).
To this end, the large discrepancy between our data and those of Gola et al. (2005) remains unresolved and cannot be explained from experimental details.
However it appears that the authors have not tested the accuracy of their
isotope ratio measurements as a function of the isotopologue mole fraction
in the presence of other species with overlapping spectra (HCl, H2O,
O3, etc.), e.g. by using a dilution series.
In any case, the strongest support for our smaller isotope effects arises
from the absence of any significant seasonal variation in the tropospheric
δ13C(CH3Cl), as shown in Sect. 3.5.
Carbon isotope modelling
Model set-up
The model used in this study is similar to previous two box models (Tans, 1997; Sapart et al., 2012; Saltzman et al., 2004; Trudinger et al.,
2004). The atmosphere is divided in two well-mixed semi-hemispheric boxes,
representing the Northern and the Southern Hemisphere, and the
interhemispheric exchange time is 360 days. The model simulates the major
sources and sinks for both the lighter (12CH3Cl) and the heavier
isotopologue (13CH3Cl), as described by Sapart et al. (2012) for
CH4. The source and sink terms from the Xiao et al. (2010) model study
serve as a starting point for our model. We use a simplified
mass balance with four source categories that has distinct isotopic source
signatures: higher plants/unknown, oceans, biomass burning and other
known sources (Sect. 3.3 and Table 4). Total net emissions were fixed at
4010 Gg a-1 with 2210 Gg a-1 in the Northern Hemisphere and 1800 Gg a-1 in the Southern Hemisphere. For each source category, the carbon
isotope source signature was randomly varied within the given uncertainties
(Table 4). Losses are specified by pseudo first-order rate coefficients. The
sinks implemented in the model (Sect. 3.4 and Table 4) are losses due to
the reaction with OH, losses to the surface ocean, losses to soils and
losses to the stratosphere. Seasonal variations were modelled with a time
step of 1 day, using monthly averaged source terms. Variations in the source
composition were modelled with a time step of 90 days, using annually
averaged source terms.
Simplified CH3Cl source and sink scheme used in the
model.
Strength (Gg a-1)
δ13C (source) / ε (sink) (‰)
Sources
Best
Range
Best
Range (1σ)
Biomass burning
910
655–1125
-47
-40 to -52
Oceans
335
210–480
-36
-30 to -42
Higher plants/unknown
2400
0–3095
-83
-70 to -96
Other known sources
365
79–1016
-45
-39 to -51
sinks
OH, Cl
3614
3564–3000
-11.2/-59
Soils∗
250
200–1000
-37
-46 to -2
Stratosphere
146
0
∗ The apparent isotope effect of the soil sink depends on
its strength. See text for more details.
Mixing ratios and isotopic composition of tropospheric
CH3Cl
Tropospheric CH3Cl has a mean global mixing ratio of about 540 pmol mol-1 (Monzka et al., 2010; Carpenter et al., 2014) and shows a
pronounced seasonal cycle with an amplitude of 85 pmol mol-1 at
Northern Hemispheric midlatitudes, (Prinn et al., 2000; Yoshida et
al., 2006), reflecting the seasonality in the OH sink (Fig. 2, upper
panel).
(a) Comparison of modelled northern hemispheric mixing
ratios (blue open diamonds) with measured mixing ratios at Mace Head, Ireland
(red crosses) for the period from 2004 to 2014 (Prinn et al., 2000). Error
bars indicate the variations in the monthly means on the 1σ level.
(b) Modelled seasonal fluctuations in the δ13C of
northern hemispheric CH3Cl using an ε of
-11.2 ‰ (red filled diamonds) and an ε of
-59 ‰ (blue filled dots) as reported by Gola et al. (2005). The
panel shows monthly averages from a 10-year simulation with seasonal
variations of up to ±10 ‰ in the combined isotopic source
signature. Error bars indicate the variations in the monthly means on the
1σ level. The grey shaded area shows reported seasonal variations
(seasonal mean ±1σ) from Thompson et
al. (2002).
Reported mean δ13C values of tropospheric CH3Cl
range from (-36.2±1.9) ‰ to (-40.8±3.0) ‰
(Tsunogai et al., 1999; Thompson et al., 2002; Redeker et al., 2007; Bahlmann
et al., 2011; Weinberg et al., 2014), suggesting an overall mean
δ13C(CH3Cl) of (-37.1±2.7) ‰. In a year-round
study carried out in Alert, Canada, Thompson et al. (2002) found no seasonal
trend in the tropospheric δ13C(CH3Cl) and no clear
correlation between the CH3Cl mixing and isotope ratios (Fig. 2,
lower panel). From their data the authors estimated the ε value
for the OH sink at less than -5 ‰. Including samples from
Frazer Point (Canada), Vancouver (Canada), Houston (Texas, USA) and Barring Head (New
Zealand), this study further revealed no indication of a latitudinal trend in
tropospheric δ13C(CH3Cl). The lack of a significant
covariation between the mixing ratios and carbon isotope ratios was
confirmed in a second year-round study (Redeker et al., 2007).
Sources
The seasonal source terms are specified for each hemisphere using monthly
means as depicted in Fig. 3. The ocean is treated as a net source for
CH3Cl with an annual net emission of 335 Gg a-1 (range: 80
to 610 Gg a-1) (Hu et al., 2013). To account for the bidirectional
nature of the gas exchange across the air–sea interface, net fluxes are
broken down into unidirectional gross uptake and emission fluxes, with the
uptake carrying the isotopic composition of the atmosphere and the emission
carrying the isotopic information of the concurrent formation and degradation
processes in the ocean. The gross uptake is calculated using an average
transfer velocity of 17 cm h-1 for CO2 (Wanninkhof, 2014)
and a mean tropospheric mixing ratio of 540 pmol mol-1. Gross
emissions are then calculated as the difference between net emissions and
gross uptake fluxes. The reader should note that this approach differs from
that of Hu et al. (2013) and results in larger gross fluxes because gross
fluxes are calculated for the entire ocean surface. Keppler et al. (2005)
estimated average isotopic composition of dissolved CH3Cl to
(-36±4) ‰. This value refers to Komatsu et al. (2004), who
reported a mean δ13C of -38 ‰ for CH3Cl
in coastally influenced waters off Japan and more enriched
δ13C values in the range of -12 ‰ to
-30 ‰ from the open north-eastern Pacific. We obtained average
δ13C values of -43±3 ‰ from a productive lagoon
in southern Portugal (Weinberg et al., 2015). Taking the biotic and abiotic
degradation of CH3Cl into account, we estimate the mean isotopic
source signature of the ocean source to be (-36±6) ‰.
Chloromethane emissions in Gg month-1 for the
Northern (a, c, e, g) and Southern Hemisphere (b, d, f, h).
(a, b) Combined emissions from higher plants and the unknown source,
(c, d) biomass burning, (e, f) ocean net emission fluxes,
and (g, h) total emissions. The emissions from the other known
sources are constant over time, with 22.8 Gg month-1 in the Northern
Hemisphere and 7.6 Gg month-1 in the Southern
Hemisphere.
We applied a source strength of 910 Gg a-1 for biomass burning (range
from 660 to 1230 Gg a-1) with 68 %, originating from the Northern
Hemisphere, and the emissions peaking during hemispheric spring (Xiao et al.,
2010). CH3Cl from biomass burning shows a δ13C of
(-47±7) ‰ (Czapiewski et al., 2002; Thompson et al., 2002).
The category “other known sources” comprises fungi wetlands and
anthropogenic emissions with a total source strength of 365 Gg a-1
(range: 79 to 1016 Gg a-1) and an averaged isotopic source signature
of (-45.5±5.5) ‰ calculated from the source signatures given
by Keppler et al. (2005). The emissions from the other known sources were
constant over time with 274 Gg a-1 in the Northern Hemisphere and
91 Gg a-1 in the Southern Hemisphere.
The source category “higher plants/missing” (900 to 3095 Gg a-1)
mainly represents emissions from the tropical rainforest (900 to
2650 Gg a-1) with minor contributions from salt marshes (80 to
160 Gg a-1), rice paddies (5 Gg a-1) and mangroves (∼50 Gg a-1). These emissions are almost equally distributed between
the hemispheres and show a slight seasonal peak during hemispheric summer.
In order to evaluate the emissions from higher plants, these emissions were
divided into two fractions by introducing a split factor. The first fraction
represents “true” emissions from higher plants, having an exceptionally
depleted isotopic source signature of (-83±15) ‰ (Saito and
Yokouchi, 2008; Saito et al., 2008). The second fraction represents an
unknown or missing source. The δ13C of this source is scaled
to match the δ13C of tropospheric CH3Cl. A more
depleted δ13C for this source would point towards additional
contributions from a lighter source, such as senescent leaf litter, whereas a
more enriched δ13C for this source points towards additional
contributions from a more enriched source.
Saito et al. (2013) recently reported on the bidirectional exchange of
CH3Cl across the leaves of tropical plants with gross uptake rates
being roughly a sixth of gross emission rates. The authors hypothesized
that the gross uptake may be related to endosymbiotic bacteria. This uptake
might affect the isotopic composition of CH3Cl from tropical
rainforests. However, because the incubation methods used in this study were
the same as those previously used to determine the isotopic composition of
CH3Cl emitted from tropical plants (Saito and Yokouchi, 2008; Saito et
al., 2008), we can reasonably assume that any potential isotopic effect of
this bidirectional exchange is included in the previously reported carbon
isotope ratios.
Sinks
The reaction with OH constitutes the single largest sink for CH3Cl,
accounting for approximately 80 % of its removal from the troposphere. For
this study, we used the OH-concentration fields from Spivakovsky et al. (2000) along with reaction rate constants of Burkholder et al. (2015) to
derive monthly resolved lifetimes for both hemispheres. The monthly loss
rates were then forced to reproduce seasonal variations of the mixing ratios
at Mace Head in the Northern Hemisphere and at Cape Grim, Tasmania,
in the Southern Hemisphere (Prinn et al., 2000). This resulted in a total
tropospheric sink (OH + Cl) of 3614 Gg a-1, which is comparable to
previous modelling studies (Xiao et al., 2010).
In most global budgets, soils are treated as a small sink for chloromethane
of about ∼250 Gg a-1, though a larger uptake exceeding
1000 Gg a-1 has been suggested (Keppler et al., 2005; Carpenter et al.,
2014). Based on Xiao et al. (2010), we a priori assumed a soil sink of 250 Gg a-1 with northern and southern hemispheric fractions of 180 and
70 Gg a-1, reflecting the interhemispheric distribution of
the land masses.
The microbial degradation of CH3Cl in soils is assigned with a
large carbon isotope effect of -47 ‰ (Miller et al., 2001, 2004).
The only study we are aware of (Redeker and Kalin, 2012) that investigates
the isotopic composition of soil derived CH3Cl reports a
δ13C of (-34±14) ‰. The soil uptake of
CH3Cl can be regarded as a coupled diffusion reaction process,
where CH3Cl is first transported into the soil and then undergoes
microbial degradation. The apparent isotope effect of such coupled processes
will depend on the isotope effects of both steps and can be estimated from
diffusion reaction models (Farquhar et al., 1982):
εapp=εd+ϵm-εd×(md-mm)md,
with εd and εm being the kinetic isotope
effects assigned to microbial degradation (47 ‰) and
diffusion (4 ‰), and where md is the
total mass of CH3Cl that enters the soil via diffusion and mm
represents the net soil sink.
The gross uptake flux (md) was estimated using a simple transfer
resistance model along with the biomes and respective active seasons as
previously employed by Shorter et al. (1995). We used an overall atmospheric
transfer resistance governing the transport to the soil surface (aerodynamic
transport resistance, quasi-laminar sublayer resistance and in canopy
transfer resistance) of 4 s cm-1, regardless of the biome, that was
derived from reported typical transfer resistances for different biomes
(Zhang et al., 2003). The soil uptake is governed by molecular diffusion
through the air-filled pore space. The soil-side transfer resistance can be
estimated from the effective diffusion in the soil column. For a first rough
estimate of the soil transfer resistance, we assume an air-filled pore space
of 0.3 (V/V) and a microbial inactive soil layer of 0.5 cm at the soil
surface. Using the Penman model (Penman, 1940) and a diffusion coefficient of
0.144 cm2 s-1 in air, we obtain a soil transfer resistance of
17 s cm-1. With a globally averaged transfer resistance of
21 s cm-1 and a CH3Cl background concentration of
540 pmol mol-1 and the land use categories from Shorter et al. (1995),
we obtain an upper limit of 1300 Gg a-1 for md.
As depicted in Fig. 4, the apparent isotope effect of the soil uptake is
bracketed by the isotope effects of both steps and decreases when increasing
the net soil uptake. The microbial degradation is rate limiting at low net
uptake rates, and the apparent isotope effect of the soil uptake is close to
that for microbial degradation. For instance, a soil sink of 250 Gg a-1
reveals an apparent ε value of -38 ‰. When
the entire chloromethane diffusing into soils is microbially degraded,
diffusion becomes the rate-limiting step, and the apparent ε
value matches that of diffusion.
In turn the imprint on the tropospheric δ13C shows a parabolic
distribution with a maximum at mm=0.5md. The isolated effect
of the soil sink would result in a maximum enrichment of 3.8 ‰ in the tropospheric δ13C that reduces to
2.1 ‰ when accounting for the concurrent reduction in
the OH sink. In this case, increasing the soil sink could even lead to a
decrease in the overall sink isotope effect once the apparent isotope effect
of the soil sink becomes smaller than the isotope effect of the OH sink.
Panel (a) shows the apparent ε of the soil
sink versus the sink strength. Panel (b) shows the resulting effect
on tropospheric δ13C. Crosses show the pure effect, e.g. in the
absence of any other fractionating sink. Open circles show the resulting
effect with an ε of -11.2 ‰ assigned to the OH sink.
Total losses were assumed to be fixed at 4010 Gg a-1.
Seasonal variations in the δ13C of tropospheric
CH3Cl
The OH-driven seasonal cycle in the tropospheric mixing ratios of CH3Cl
implies an inverse covariation in the δ13C of tropospheric
CH3Cl to an extent that is closely linked to the kinetic isotope effect
of the OH sink.
Our model resembles mean tropospheric mixing ratios of about
540 pmol mol-1 (Monzka et al., 2010; Carpenter et al., 2014) and the
seasonal cycles of CH3Cl in both hemispheres within ±4 %
(Fig. 2 upper panel). In our simulations, an ε value of
-59 ‰ for the OH sink produces an inverse covariation of the
δ13C(CH3Cl) with the CH3Cl mixing ratios with a
seasonal amplitude of 9.2 ‰, whereas our new smaller ε
value of -11.2 ‰ results in a seasonal amplitude of only
1.7 ‰ (Fig. 2 lower panel), fitting quite well to the measured
variation given by Thompson et al. (2002).
Random variations of ±10 ‰ in the isotopic source signatures,
seasonal variations of the emission functions and variations in the soil sink
resulted in seasonal fluctuations of up to ±10 ‰ in the
combined isotopic source signature (see Supplement for further details). As
already noted by Tans (1997) the large tropospheric background strongly
attenuates temporal variations. In our model simulations these seasonal
variations in the combined isotopic source signature resulted in systematic
seasonal variations of the northern hemispheric δ13C of less
than ±0.7 ‰ attributable to the isotopic source signal and a
scatter of up to ±1.0 ‰ in the monthly mean
δ13C(CH3Cl) values in the Northern Hemisphere. In all our
model simulations these variations in the combined isotopic source did not
significantly affect the differences in the seasonal amplitude of the
tropospheric δ13C signal. They may have an imprint on the
tropospheric δ13C(CH3Cl) signal when applying an
ε value of -11.2 ‰ to the OH sink. However, they are
largely obscured when applying an ε value of -59 ‰
to the OH sink (Fig. S6 in the Supplement). Masking the isotope effect of
this large ε value would require seasonal variations of about
50 ‰ in the combined source signatures that are inversely correlated to
the OH sink. In sum, the lacking covariation between the mixing ratios and
carbon isotope ratios strongly supports our new ε value of
-11.2 ‰ and makes the previously reported larger ε
value highly unlikely.
Implications for the tropical rainforest source
An ε value of -59 ‰ for the OH sink
requires a mean mass-weighted isotopic source composition of -84.5 ‰ to balance the tropospheric δ13C(CH3Cl) of (-36.4±2.1) ‰ (Thompson
et al., 2002), as shown in previous studies. Apart from large emissions from
higher plants (Keppler et al., 2005; Saito and Yokouchi, 2008) in
tropical rainforests, this isotope effect suggests additional substantial
emissions from an even more depleted source, such as senescent leaf litter
(Keppler et al., 2005; Saito and Yokouchi, 2008). In contrast, the revised
smaller ε value of -11.2 ‰ requires a mean
isotopic source signature of -48.5 ‰, which is close to
the mass-weighted δ13C of all other known sources excluding
higher plants. Along with higher plant emissions of 2200 Gg a-1, the
new ε value of -11.2 ‰ yields a mean
tropospheric δ13C(CH3Cl) of -56 ‰,
which is depleted by almost 20 ‰ in comparison to the mean
reported tropospheric δ13C(CH3Cl).
We performed more than 10 000 steady-state runs with random variations in
the isotopic composition of tropospheric CH3Cl (-36.4±2.1 ‰), the isotopic source signatures (Table 4) and the
isotope effect of the soil sink to assess the range of CH3Cl emissions
from higher plants. The source category “higher plants” was divided in two
fractions, one representing “true” emissions from higher plants and the
other representing missing emissions. The δ13C of the missing
emissions was always scaled to match the tropospheric δ13C(CH3Cl).
As shown in Fig. 5, the strength of the tropical rainforest source is
directly linked to the strength and isotopic composition of missing
emissions. A tropical rainforest source of (600±200) Gg a-1
suggests missing emissions of (1600±200) Gg a-1, requiring a
δ13C of (-45±6) ‰ to balance the
tropospheric δ13C(CH3Cl). This δ13C is close
to the mean isotopic composition of all other known sources. Increasing the
emissions from all other known sources within the given ranges might reduce
the missing emissions by about 500 Gg a-1. A further increase of the
tropical rainforest emissions results in an equivalent reduction in missing
emissions but requires a more enriched δ13C for the missing
emissions. For instance, balancing a tropical rainforest source of (1100±200) Gg a-1 requires missing emissions of the same magnitude
with a δ13C of (-31±6) ‰. This
is at the upper end of reported source signatures and may thus serve as a
boundary with which to constrain the rainforest source from an isotopic perspective.
Modelled isotopic composition of the missing source versus tropical
rainforest emissions on the lower x axis and missing emissions on the upper
x axis (rainforest = 2200 – unknown). The black line shows the best
estimate derived from the mean isotopic source signatures. Orange dots
indicate the range uncertainty (1σ) from uncertainties (1σ)
in the δ13C of biomass burning (±7 ‰), ocean net
emissions (±6 ‰) and other known sources (±6 ‰).
Yellow dots mark the additional uncertainty from the δ13C of
the tropical rainforest source (±10 ‰). The green column
indicates the carbon-density-based estimate of the rainforest source and the
grey bar indicates the range in δ13C of biomass burning and
the mean from all sources excluding the tropical rainforest.
Carbon-density-based revision of the tropical rainforest source
Interestingly, support for our lower estimate arises from previous studies
on the tropical rainforest CH3Cl source when using above-ground carbon
density instead of coverage area for upscaling the CH3Cl emission
factors.
Large uncertainties in upscaling of locally derived plant emissions to global
scales can arise from (i) temporal variations in the emissions and (ii) spatial
variability in environmental drivers, species composition and vegetation
cover. Within the widely used FAO land cover classes, forests are defined as
land with tree cover exceeding 10 %, a potential tree height of 5 m
and an area of at least 0.5 ha (FAO, 2012). Tropical rainforests encompass
such sparsely covered areas with a carbon density of only a few Mg ha-1
(Asner et al., 2010; Pereira et al., 2016) as well as very dense mature
rainforests with a canopy height of more than 40 m and above-ground carbon
densities sometimes exceeding 300 Mg ha-1 (Kato et al., 1978). This
suggests a large variability in biomass that cannot be assessed with the
previously used area-based upscaling approaches. Area-based estimates may be
improved by leaf area index or above-ground carbon-density-based approaches.
There is some indication that CH3Cl is mainly emitted by mature
trees (Saito and Yokouchi, 2008; Saito et al., 2008, 2013). This is more
readily reflected by carbon density than by leaf area. Further, the available
carbon density data products allow direct discrimination between tropical
forests and other tropical vegetation. We thus propose a carbon-density-based
upscaling approach of experimentally derived emission factors to reduce
uncertainties arising from the spatial variability in above-ground biomass.
We first convert reported area-based emission factors to carbon-density-based
emission factors and then multiply them with the carbon stock of the tropical
rainforest:
F(CH3Cl)=EF×CRFCst.
Here, F(CH3Cl) is the source strength [Gg a-1], EF is the
experimentally derived emission factor [Gg ha-1 a-1], Cst is
the above-ground carbon density assigned to the sampling site [Gg ha-1]
and CRF is the estimated total above-ground biomass [Gg] of the
respective biome, in this case the tropical rainforest.
Calculation of carbon-density-based CH3Cl emission factors
from previously reported area-based emission factors.
Site
Method
Carbon density
Emission per
Reference
leaf dry mass
area
carbon density
Mg ha-1
µg g-1 h-1
µg m-2 h-1
g ha-2 a-1
g Mg-1 a-1
Glass house, Japan
branch enclosure
325
0.32
74.0
Yokouchi et al. (2000)
Pasoh Forest Reserve,
micrometeorological
265
14.0
1226
4.6
Saito et al. (2013)
Malaysia
branch enclosure
0.03
5.0
Pasoh Forest Reserve,
branch enclosure
265
0.026
7.0
615
2.3
Saito et al. (2013)
Malaysia
Danum Valley,
branch enclosure
265
0.03
12.0
1051
4.0
Blei et al. (2010)
Borneo, Malaysia
Surinam, French
gradient above canopy
160
9.5
832
5.2
Gebhardt et al. (2008)
Guyana
Mean
931
4.0
SD
265
1.2
The first direct evidence for strong CH3Cl emissions from tropical
plants came from branch incubations of tropical plants in a greenhouse
(Yokouchi et al., 2000). This study revealed particularly high emission from
dipterocarp species that are dominant in tropical lowland rainforests of
southern and south-eastern Asia and suggested mean CH3Cl emissions of 74 µg m2 h-1. Several follow-up studies carried out in
tropical rainforests reported 10- to 5-fold lower fluxes (Saito et al.,
2008, 2013; Gebhardt et al., 2008; Blei et al., 2010). We exclude the high-emission factor from the greenhouse study from our reanalysis of the
tropical rainforest source and focus on the studies, providing
experimentally derived emission factors for CH3Cl emissions from
tropical forests and allowing for a sufficient estimate of carbon densities
assignable to them. Details on these studies are provided in Table 5. Three
studies have been carried out in lowland tropical rainforests of
south-eastern Asia, and one has been carried out over Surinam in South America. We are not
aware of any CH3Cl flux measurements from African tropical rainforests.
Two studies relied on branch or leaf incubation to measure CH3Cl fluxes
(Saito et al., 2013; Blei et al., 2010). A third study used a
micrometeorological approach and in addition performed leaf and branch
incubation (Saito et al., 2008). The remaining study (Gebhardt et
al., 2008) derived CH3Cl emissions factors from concentration gradients
above the rainforest. The concentration gradients were obtained from
canister samples taken at different heights above the rainforest from an
airplane. The results from branch and leaf incubations were first normalized
to leaf dry weight and then converted to area-based emission factors using
reported allometric data for south-eastern Asian tropical lowland rainforests
along with assumptions of the distribution and abundance of the investigated
species. The mean area normalized fluxes (8.0 µg m2 h-1±45 %) from these studies show a notably larger
variability than the original leaf biomass normalized fluxes (0.028 µg g-1 h-1±10 %), although all three studies referred
to the same allometric data (Yamakura et al., 1986) in their conversion.
Note that the study reporting the lowest emissions factors from branch
enclosures reported almost 3 times higher fluxes using a
micrometeorological approach. In sum, the area-based factors agree within a
factor of 3 and range from 5.0 to 14 µg m2 h-1
(9.1 µg m2 h-1±37 %).
Comparison of calculated area-based and carbon-density-based
emissions from tropical forests.
Region
Area
C density
Area-based
Carbon-density-based
Percentage of area
106 ha
106 g ha-1
Gg a-1
Gg a-1
based
Brazil
586
1121
496±143
247±73
50
Indonesia
165
1121
140±40
69±20
50
Congo
205
921
173±50
71±21
41
Tropical Africa
775
621
655±189
182±53
28
Tropical America
1209
771
1022±295
356±103
35
Tropical Asia
474
981
401±116
176±51
44
Pantropics
2458
781
2079±601
720±212
35
Tropical Africa
393
822
332±96
121±36
36
Tropical America
788
1162
666±192
343±101
52
Tropical Asia
289
1192
244±71
129±38
53
Pantropics
1470
1052
1243±359
580±171
47
1 Saatchi et al. (2011); 2 Baccini et al. (2012).
The three south-eastern Asian studies refer to a dense and mature dipterocarp
forest with an above-ground carbon density of (265±44) Mg C ha-1 (Yamakura et al., 1986) that we apply here. For the
study carried out above the rainforest of Surinam, we derived a carbon
density of (160±15) Mg ha-1 from carbon density maps (Saatchi et
al., 2011; Baccini et al., 2012).This range agrees with the FAO estimate for
French Guyana (FAO, 2015) and is supported by several field surveys carried
out in this region (Chave et al., 2001, 2008; Saatchi et al., 2007). With
this, we obtain a mean carbon-density-based emission factor of (4.0±1.2) g Mg-1, referring to a mean carbon density of
202 Mg ha-1. This is well above the average tropical rainforest carbon
density, ranging from 96 to 117 Mg ha-1 (Baccini et al., 2012; Saatchi
et al., 2011; Köhl et al., 2015; FAO, 2015).
In consequence, our carbon-density-based estimates of the tropical
CH3Cl source are 30 % to 70 % lower than the respective area-based
estimates (Table 6). The difference is in the range of 30 % for dense old
grown evergreen forests such as the Tierra Firme forests of French Guyana,
between 40 % and 50 % for the moist tropical rainforest, and increases
to almost 70 % for the entire pantropical forests including dry tropical
forests, degraded forests and plantations. This trend reflects the
decreasing trend in carbon density in the tropical rainforest biomes as well
as the effect of forest degradation. Regardless of the source for the carbon
density estimates, our approach suggests a tropical rainforest CH3Cl
source of (670±250) Gg a-1, which is 53 % to 65 % lower than
the respective area-based estimates in the range of 1200 to 2000 Gg a-1 (Saito et al., 2008, 2013; Gebhardt et al., 2008; Blei et al., 2010).
The uncertainty in the area-based emission factors is estimated to 24 %
from the standard deviation of the reported means. Additional uncertainties
for our carbon-density-based upscaling (compared to the previous area-based upscaling) arise from the uncertainties in the total above-ground
carbon stocks (±8.6 %) and the site-specific carbon density
(±15 %). Using error propagation, we estimate the total
uncertainty of our approach to be ±30.4 %. However, we note an urgent
need for more detailed flux studies. Currently there is no information about
how physiological and environmental drivers might affect CH3Cl
emissions from tropical rainforests. Apart from the observation that some
members of the Dipterocarpaceae family are particularly strong emitters of
CH3Cl, this also holds true with respect to species composition.
Conclusions
We reported new ε values for the reaction of CH3Cl with OH
and Cl of (-11.2±0.8) ‰ (n=3) and (-10.2±0.5) ‰, which are 5 to 7 times
smaller than the previously reported ε values for these
reactions. Strong support for the reliability of our new fractionation
factors arises from the absence of any significant covariation in the
mixing and carbon isotope ratios of tropospheric CH3Cl.
Conjoining our new KIEs of the tropospheric CH3Cl sinks and the
biomass-based upscaling of previously reported emission factors suggests a tropical
vegetation source of only (670±200) Gg a-1, which is about
3-fold smaller than suggested in current budgets. We assign δ13C of -45±6 ‰ to the missing emissions of
(1530±200) Gg a-1. Notably increasing the soil sink by 750 Gg a-1 and decreasing biomass burning emissions by 460 Gg a-1, as
suggested in the latest assessment on ozone-depleting substances (Carpenter
et al., 2014), would substantially increase this gap but have a negligible
effect on the isotopic composition of the missing emissions. The δ13C value of the missing emissions matches with the mean source
signature of the other known sources (except rainforests). Increasing these
emissions within the given ranges might reduce the gap to (1100±200) Gg a-1.
From a purely isotopic perspective, in particular larger
emissions from biomass burning could further reduce this gap. However, this
is highly speculative as virtually any source combination providing a mean
δ13C of -45±6 ‰ could fill the gap.
With CH3Cl being the single largest natural carrier of chlorine to the
stratosphere, predicting future baselines of stratospheric chlorine requires
a better understanding of the global CH3Cl cycle and an identification
of the missing emissions.