Introduction
Nitrous oxide (N2O) is a potent greenhouse gas and stratospheric ozone-depleting substance emitted from various sources on the Earth surface. It is
injected into the stratosphere in tropical upwelling regions. In the
stratosphere, it is transported meridionally by the Brewer–Dobson
circulation and is decomposed by photolysis (Eq. 1) and photooxidation
(Eq. 2) with approximate shares of 90 and 10 %, respectively (Minschwaner
et al., 1993).
N2O+hν→N2+O(1D)
N2O+O(1D)→2NO→N2+O2
Mixing ratios of trace gases such as N2O and their interrelationships
are regarded as useful tools to establish a detailed picture of stratospheric
circulation (Plumb, 2007), and previous observations showed compact tracer
relationships that depend on the latitude (e.g., Michelsen et al., 1998).
Natural abundance ratios of N2O isotopocules, molecular species that
only differ in either the number or position of isotopic substitutions
(Coplen, 2011), are useful tracers for elucidating the sources and
physicochemical records of N2O because isotopocule ratios reflect the
isotopic compositions of source materials; isotope effects specific to each
chemical and physical process relevant to formation, transport, and
decomposition of N2O; and mixing of various air and water masses (Toyoda
et al., 2017). In the context of stratospheric distribution, isotopocule
ratios are unique in their ability to provide the degree of photochemical
decomposition and the relative importance of the above-mentioned two
decomposition pathways (Toyoda et al., 2001; Röckmann et al., 2001).
Stratospheric distributions of N2O isotopocules have been studied using
scientific balloons and aircraft. For balloon observations, (a) whole air
samples are collected using cryogenic samplers and are evaluated using
isotope ratio monitoring mass spectrometry (IRMS) (Kim and Craig, 1993; Rahn
and Wahlen, 1997; Röckmann et al., 2001; Toyoda et al., 2001, 2004; Kaiser et al., 2006); alternatively, (b) remote measurements are
conducted using Fourier-transform infrared spectrometry (Griffith et al.,
2000). Such observations have revealed vertical profiles up to 35 km
altitude. For aircraft observations, whole air samples are collected using
pressurizing pumps, with subsequent evaluation by IRMS (Rahn and Wahlen,
1997; Park et al., 2004; Kaiser et al., 2006). Although these observations
are limited to up to ca. 20 km altitude, mean vertical profiles can be
obtained from horizontally extended areas such as Arctic polar vortexes.
Earlier studies showed that a decrease in the mixing ratio of N2O with
altitude is accompanied by enrichment of isotopocules of N2O heavier
than the major one (14N14N16O) (Kim and Craig, 1993; Rahn and
Wahlen, 1997). The results were consistent with the expected isotopocule
fractionation during photolysis (Yung and Miller, 1997), but the apparent
degrees of fractionation (isotopic fractionation, ε) differed
between those of the lower (< 20–25 km where N2O mixing ratio
> 180–250 nmol mol-1) and middle (> 20–25 km or
< 180–250 nmol mol-1) stratosphere (Toyoda et al., 2001;
Röckmann et al., 2001; Kaiser et al., 2006; Park et al., 2004). They are
always smaller than fractionations obtained from laboratory photolysis
experiments (Röckmann et al., 2000; Turatti et al., 2000; Zhang et al.,
2000; Kaiser et al., 2002b, 2003) or absorption spectra (von Hessberg et al.,
2004). Moreover, the values of ε for the middle stratosphere
were found to depend on the latitude and season (Toyoda et al., 2004; Kaiser
et al., 2006; Park et al., 2004). Such variation has been regarded as a
result of photochemical and transport processes (McLinden et al., 2003; Park
et al., 2004; Kaiser et al., 2006), but it has not been fully examined
because of the lack of measurements taken over the tropical upwelling
regions. As noted above, the tropical stratosphere is the starting point of
the meridional transport of N2O injected from the troposphere and the
rates of photochemical reactions are faster than those in the extratropics
because of stronger actinic flux. Although there are a few reports on
vertical profiles of N2O isotopocules over India (18∘ N)
(Röckmann et al., 2001; Kaiser et al., 2006), the isotopic composition of
N2O in upwelling tropical air has not been characterized.
Map showing balloon launching sites. EQP, eastern equatorial
Pacific; BIK, Biak Island, Indonesia.
Another controversial problem about the stratospheric N2O is whether the
photooxidation sink (Eq. 2) has a larger contribution than 10 % in the
lower stratosphere. Kaiser et al. (2002a) found that the ratios of
fractionations of isotopocules during photolysis and photooxidation are
distinct and they estimated that a much larger
fraction (up to 100 %) is removed by photooxidation at least in the lower
stratosphere (N2O mixing ratios > 300 nmol mol-1) (Kaiser et
al., 2006). However, similar but a little simplified analyses by Park et
al. (2004) and Toyoda et al. (2004) (see Sect. 3.4) could not detect
significant differences in the isotopic fractionation ratios between the
lower and middle stratosphere, and Park et al. (2004) speculated that the
isotopic fractionation ratios could be affected not only by the relative
share of the two sink pathways but also by other factors such as transport.
This study was conducted to ascertain the vertical profiles of N2O and
its isotopocules over the Equator and to examine the factors that control
the apparent isotopic fractionation in the stratosphere by comparing results
from earlier studies and three-dimensional chemical transport model
simulation.
Experiments and model simulation
Whole air sampling over the Equator
Stratospheric air samples were collected using balloon-borne compact
cryogenic samplers (J-T samplers) (Morimoto et al., 2009). The sampler
consists of an evacuated 800 cm3 stainless steel sample flask (SUS304),
a cooling device called a Joule–Thomson (J-T) mini cooler, a 2 L
high-pressure neon gas cylinder, pneumatic valves, solenoid valves, a
100 cm3 high-pressure N2 gas cylinder for actuation of pneumatic
valves, an electronic controller with a GPS receiver, a telemetry
transmitter, and batteries. The J-T mini cooler can produce liquid neon from
high-pressure neon gas that is pre-cooled by liquid nitrogen. The liquid neon
is used as refrigerant to solidify or liquefy the stratospheric air.
Contrasted against the larger sampling system used in our previous
observations, which was about 250 kg with 12 sample flasks in a Dewar flask
filled with liquid helium, the J-T sampler weighs ca. 20 kg, with
operational and logistic advantages at remote sites such as remote islands or
polar regions.
Sampling over the eastern equatorial Pacific (0∘ N,
105–115∘ W) was conducted on 4, 5, 7, and 8 February 2012 during
the KH-12-1 cruise of R/V Hakuho-maru, JAMSTEC as a part of the
Equatorial Pacific Ocean and Stratospheric/Tropospheric Atmosphere Study
Program. For each balloon flight, a 5–8 L volume STP of air sample was
collected by a single sampler at programmed altitude between 19 and 29 km.
The sampler then descended by parachute. It was later recovered on the sea.
Another sampling campaign was conducted at Biak Island, Indonesia
(1∘ S, 136∘ E), on 22, 24, 26, and 28 February 2015 as part
of a small size project by ISAS/JAXA (Hasebe et al., 2018). For each balloon
flight, two samplers integrated into a single gondola were launched from the
observatory of the National Institute of Aeronautics and Space of the
Republic of Indonesia (LAPAN). Samples were collected at two altitudes. In
total, we obtained seven samples: two samples on each of four flights, with
one sampling failed.
Locations of launching sites are shown in Fig. 1. Balloon trajectories are
portrayed in Fig. S2 in the Supplement. Sampling was conducted while the
balloon was ascending, except for the flight on 5 February 2012. Typical altitude
range was about 2 km, and we took the central value of the range as sampling
altitude.
Analysis of mixing ratio and isotopocule ratios
At Tohoku University, the mixing ratio of N2O was measured using gas
chromatography with electron capture detection (GC-ECD) with precision of
1 nmol mol-1 (Ishijima et al., 2001). The isotopocule ratios, defined
as follows, were measured at Tokyo Institute of Technology using gas
chromatography – isotope ratio mass spectrometry (Toyoda et al., 2004;
Toyoda and Yoshida, 2016).
δX=(Rsample-Rstandard)/Rstandard
Therein, X denotes 15Nα, 15Nβ, or 18O,
and where R denotes 14N15N16O/14N14N16O,
15N14N16O/14N14N16O, or
14N14N18O/14N14N16O of the sample and standards
(Toyoda and Yoshida, 1999). The δ value is expressed as the per mill
(‰) deviation relative to atmospheric N2, and Vienna Standard
Mean Ocean Water (VSMOW), respectively, for nitrogen and oxygen. In addition
to δ15Nα and δ15Nβ, the δ
value for bulk N and 15N-site preference (SP) are often used as
illustrative parameters:
δ15Nbulk=(δ15Nα+δ15Nβ)/2,SP=δ15Nα-δ15Nβ.
Duplicate analyses were made for a set of two runs: monitoring of molecular
ion for determination of δ15Nbulk and δ18O
and NO+ fragment ion for determination of δ15Nα. A
300–400 cm3 STP aliquot of the sample air was introduced into the
analytical system from the sample flask in a single run. Typical precisions
of the isotopic analyses are < 0.1 ‰ for
δ15Nbulk, < 0.2 ‰ for δ18O, and
< 0.5 ‰ for δ15Nα, although they were
slightly worse for samples collected at higher altitudes because of the lower
N2O mixing ratio.
Vertical profiles of mixing ratio (a),
δ15Nbulk (b), SP (c), and
δ18O (d), of N2O observed over the Equator (pink
symbols). Previously published results obtained over Japan (black and red
symbols), Sweden (blue), and Antarctica (green) (Toyoda et al., 2001, 2004),
and India (orange; Kaiser et al., 2016; Röckmann et al., 2001) are also
shown. In the legend, launch sites and dates are shown, respectively, by
three characters and six digits in yymmdd format. SBC, Sanriku Balloon
Center, Japan; ESR, Esrange, Kiruna, Sweden; SYO, Syowa station, Antarctica;
HDB, Hyderabad, India; EQP, eastern equatorial Pacific; BIK, Biak Island,
Indonesia. See also Table S2 for details.
Correlation between mixing ratio and δ15Nbulk
of N2O (Rayleigh plot). The high mixing ratio range (> ca.
120 nmol mol-1) in (a) is enlarged in (b). Both
parameters are normalized to their values at the time when the corresponding
air mass entered the stratosphere (see Eq. 6 in the text). Grey solid and
broken lines show slopes obtained respectively from laboratory broadband
photolysis experiments (Kaiser et al., 2002b, 2003) and photooxidation
experiments (Kaiser et al., 2002a; Toyoda et al., 2004).
To analyze the relation between the N2O mixing ratio ([N2O]) and
isotopocule ratio (δ) in a Rayleigh fractionation scheme (Eq. 6),
measured values must be normalized with respect to the values before the air
mass enters the stratosphere.
(1+δ)/(1+δ0)={[N2O]/[N2O]0}ε
In Eq. (6), subscript 0 signifies a tropospheric value; ε is the
isotopic fractionation. Because the tropospheric mixing ratio and isotopocule
ratios are known to have secular trends, [N2O]0 and δ0
were estimated as follows. First, the age of the measured air mass was
estimated based on the mixing ratio of CO2 (Engel et al., 2009), which
was also measured for the same air sample. Then, the N2O mixing ratio at
the time when the air mass was in the troposphere was calculated using the
estimated age of air and the secular trend of tropospheric mixing ratio
observed by the ALE/GAGE/AGAGE project (Prinn et al., 2000). We used AGAGE
data from Mace Head (Ireland) to calculate [N2O]0 for stratospheric
air in the tropics (this study) and the Northern Hemisphere (in which our
previous observations were conducted) and those from Cape Grim (Tasmania) for
the Southern Hemisphere (our previous observation in Antarctica). For
calculating δ0, the secular trends observed at Hateruma Island,
Japan (Toyoda et al., 2013), and Cape Grim (Park et al., 2012) were used,
respectively, for the tropics/Northern Hemisphere and the Southern
Hemisphere. In Table S3 in the Supplement we compare how much this correction
regarding the age of air changed the position of each data point in Figs. 3
and 5. Typically, the term related to mixing ratio
(-ln{[N2O] / [N2O]trp}) is decreased by
0.2–3 % when we use [N2O]0, the value when the air mass actually
entered into the stratosphere, instead of the value at the same time of the
observation, [N2O]trp. The isotopic terms (ln{(1+δ)/(1+δtrp)}) are either increased or decreased depending on
their secular trends, and they are changed by 0.2–3 %.
Simulation using a three-dimensional chemical transport model
To examine the factors controlling the stratospheric distributions of
N2O isotopocules, a numerical simulation was conducted using the Center
for Climate System Research/National Institute for Environmental
Studies/Frontier Research Center for Global Change atmospheric general
circulation model with chemical reactions (CCSR/NIES/FRCGC ACTM) (Ishijima et
al., 2010, 2015). Because Ishijima et al. (2015) have
already given a detailed description of the N2O isotopocule model, we
briefly explain it here.
The N2O photolysis rate was calculated for 15 bins from 178 to 200 nm
and for three bins from 200 to 278 nm using a scheme incorporating the
parameterization of Minschwaner et al. (1993) (Akiyoshi et al., 2009) and by
a main radiation – photolysis scheme of the ACTM (Sekiguchi and Nakajima,
2008). Fractionation of N2O isotopocules was simulated using
wavelength-dependent and temperature-dependent isotopic fractionations
(ε) for 14N15N16O and 15N14N16O
reported by von Hessberg et al. (2004) although the ε for
14N14N18O was estimated from the relation between apparent
ε for each isotopocule observed in the stratosphere due to the
lack of suitable experimental reports. The model transport was nudged to
ERA-Interim reanalysis (Dee et al., 2011) for horizontal winds and
temperature at 6-hourly time intervals. Regarding the photooxidation sink of
N2O, the concentration of O(1D) was calculated online in the ACTM;
ε values were calculated as described by Kaiser et al. (2002a).
Comparison of absolute values of the isotopocule fractionation
(|ε|) for 15Nbulk, 15Nα, and
18O of N2O between observations and laboratory experiments. L and M
respectively refer to the lower and middle stratosphere with boundary mixing
ratio of about 170 nmol mol-1 (-ln{[N2O]/[N2O]0}=0.6) and 260 nmol mol-1 (-ln{[N2O]/[N2O]0}=0.2) for extratropics and tropics, respectively, based on the Rayleigh plot
shape (Fig. 3). The respective |ε| for Japan, Sweden, and
Antarctica are from Toyoda et al. (2004). Those for photolysis and
photooxidation experiments are referred from reports by Kaiser et al. (2002a, b,
2003) and Toyoda et al. (2004). Error bars show either the standard deviation
for the mean value (observation over Japan and photooxidation experiments),
standard error associated with linear regression in Rayleigh plot
(observations except Japan), or the possible range for stratospheric
conditions (photolysis experiments).
While both the surface emissions and the photolytic isotopocule
fractionations were optimized in the earlier work by Ishijima et al. (2015),
only the former was optimized in the present study. This is because we
considered that it would be better to keep the experimentally determined
original isotopocule fractionations for the purpose of comparison between the
model and the observations in the stratosphere. Moreover, we found that
apparent isotopocule fractionations obtained by the model simulation become
much closer to those by the balloon observations by replacing the
meteorological data from JRA-25 (Onogi et al., 2007) with those from
ERA-interim. This is probably because dynamics and chemical reactions in the
model were improved by the replacement of the meteorological reanalysis data
for nudging. Surface emissions of the four N2O isotopocules were
optimized in the manner described in an earlier report (Ishijima et al.,
2015), with emissions modified to reproduce observed trends (Röckmann and
Levin, 2005) and interhemispheric differences (Ishijima et al., 2007) of
atmospheric N2O isotopocule mixing ratios. Consequently, the estimated
emissions were used for a forward simulation of four N2O isotopocules in
the atmosphere from the surface to the stratosphere in this study. The
emissions and tropospheric values are reasonable (see Supplement) compared to
those of past studies (e.g., Toyoda et al., 2013, 2017) in
terms of the necessary order of precision for analysis of the large vertical
profiles in the stratosphere in this study.
Results and discussion
Vertical profiles of the N2O mixing ratio and isotopocule
ratios over the Equator
In all, 11 samples (4 from the eastern Pacific, 7 at Biak Island) were
collected at target altitudes; of them, 10 were measured for N2O
isotopocules. Figure 2 presents vertical profiles of the N2O mixing
ratio, δ15Nbulk, SP, and δ18O observed over
the Equator. Data from our previous observations over Japan, Sweden, and
Antarctica and those from observations by Röckmann et al. (2000) and
Kaiser et al. (2006) conducted over India are also shown. The height of the
tropical tropopause layer (TTL) is typically 14–18.5 km (e.g., Fueglistaler
et al., 2009), whereas the tropopause height was 12–16 km over Japan and 9
or 10 km over Sweden and Antarctica (Table S2). As observed at mid-latitudes
and high latitudes, the mixing ratio decreases with height; isotopocule
ratios increase with height over the Equator. However, the vertical gradient
is weaker at lower latitudes. Our observation over the Equator shows the
weakest gradient. Although a slight difference in mixing ratio was observed
for 20–25 km, the two equatorial profiles obtained at different longitudes
over the Equator agreed quite well. We combined the two datasets as a single
one for further examinations.
Correlations between δ15Nβ and δ15Nα of N2O (a) and between δ18O and
δ15Nbulk of N2O (b). The δ values
are normalized as noted in the text. Grey solid and broken lines show slopes
obtained respectively from laboratory broadband photolysis experiments
(Kaiser et al., 2002b, 2003) and photooxidation experiments (Kaiser et al.,
2002a; Toyoda et al., 2004).
Correlation between mixing ratio and isotopocule ratios: apparent
isotopocule fractionations
In Figs. 3 and S3–S5, the isotopocule ratios are shown against the N2O
mixing ratio after the normalization described in Eq. (6) (Rayleigh plot).
The equatorial data for lower altitudes
(-ln{[N2O] / [N2O]0}< 0.2 or [N2O] > ca.
260 nmol mol-1) are on the line defined by the data for lower
altitudes (-ln{[N2O] / [N2O]0}< 0.6 or
[N2O] > 170 nmol mol-1) over middle latitudes and high
latitudes. The linear relation is consistent with isotopocule fractionation
during the decomposition of N2O in a closed system, although the slope
of the line, which corresponds to isotopic fractionation (ε), is
markedly lower than that obtained by laboratory photolysis experiments (see
below). However, the three data points obtained at altitudes corresponding to
-ln{[N2O] / [N2O]0}> 0.2 ([N2O] < ca.
260 nmol mol-1) show systematic deviation from the line and seem to
define another line (Fig. 3b). A similar deviation or bending
structure of the Rayleigh plot has also been observed at middle to high
latitudes (Fig. 3a, from the points where the x axis value is ca. 0.5)
(Toyoda et al., 2004). We therefore compare the slope of the lines obtained
for observations at various latitudes and for laboratory simulation
experiments.
As portrayed in Fig. 4, absolute values of ε(|ε|)
for 15Nbulk, 15Nα, and 18O in the
equatorial lower stratosphere are slightly higher than those of middle
latitude and high latitude lower stratosphere, but they are still only about
half of the ε obtained by broadband photolysis experiments
(Kaiser et al., 2002b, 2003). In contrast, |ε|
in the higher region (or middle stratosphere) show larger values. They are
the largest over the Equator except for ε(18O). The
equatorial values of ε almost as large as those of photolysis.
It is also noteworthy that |ε| in the middle stratosphere in
the arctic polar vortex (Sweden) is as small as that in the lower
stratosphere and that latitudinal and year-to-year or seasonal variation are
small in the lower stratosphere compared to variations in the middle
stratosphere. Although the similar latitudinal and altitudinal dependence of
ε has been reported previously for latitudes ranging from 18 to
89∘ N (Park et al., 2004; Kaiser et al., 2006), our equatorial data
show that the change in ε occurs at altitude corresponding to
higher N2O mixing ratio. The ε value is exactly what would
be expected during the N2O photolysis as discussed below.
Cause of the variation of stratospheric ε
We now discuss causes of (1) lower |ε| value in the lower
stratosphere, (2) increase in |ε| in the middle stratosphere,
and (3) the largest |ε| in the equatorial middle stratosphere
based on two factors: photochemical and transport processes.
Photochemical processes
During photochemical decomposition of N2O, ε reportedly
depends on the wavelength that photolyzes N2O, the relative share of
photolysis and photooxidation pathways (Eqs. 1 and 2), and temperature
(Toyoda et al., 2004; Kaiser et al., 2006). Moreover, because of transport
processes the stratosphere cannot be always treated as an isolated system
which is a prerequisite for Rayleigh fractionation model. The ratio of
ε values for independent isotopocules (e.g., ε(15Nbulk)/ε(18O)), however, has been identified
as a useful parameter to distinguish photolysis and photooxidation (Kaiser et
al., 2002a) because its sensitivity to wavelength and temperature is small
and it is not affected by mixing process. Figure 5 shows the data obtained in
this study and some previous ones in δ–δ space after the
normalization described in Eq. (6). Especially in Fig. 5b, almost all data
show a compact linear relation without bending or curved structure apparent
in the Rayleigh diagram. The slope, which corresponds to the ratio of
ε values, is very close to the one expected for photolysis. This
agrees with the fact that photochemical decomposition of N2O is mainly
caused by photolysis (Minshwaner et al., 1993), although the small
fluctuation in the lower left region in Fig. 5a will be discussed later.
Transport processes
Transport processes accompanied by mixing of variously aged stratospheric air
has been considered as the major cause of lower |ε| value in
the stratosphere than in the laboratory photochemical decomposition (Park et
al., 2004; Kaiser et al., 2006). Our new observation revealed that all the
N2O isotopocules are fractionated by the almost ideal Rayleigh process
in the middle stratosphere over the deep tropics, where the stratosphere is
effectively the most isolated relative to all other regions. This underlines
how much transport and mixing affect the apparent ε value.
We then consider the effect of transport on the apparent ε at
different latitudes with a conceptual two-dimensional circulation model in
the tropical and extratropical stratosphere that was proposed to explain
tracer–tracer correlation (Plumb, 2002). In the tropics, N2O is
decomposed gradually during upwelling of the air mass injected from the
troposphere. The uppermost tropical air mass X0 is then transported to
middle latitudes and higher latitudes, where it begins downwelling. Because
the vertical ascent rate in the tropics is faster than quasi-horizontal
transport out to the extratropics and much faster than the quasi-horizontal
transport of extratropical air into the tropics, there is an apparent
transport barrier between the tropics and extratropics (Plumb, 2007).
Nevertheless, entrainment of air mass Yi across the subtropical edge
separating the two regions must occur to compensate for mass flux in the lower
region (Fig. 6).
Conceptual two-dimensional circulation model to analyze mixing
processes between tropics and extratropics (from Plumb, 2002). The X0 is
the uppermost tropical stratospheric air mass; Xi (i=1–8) are air
masses formed by mixing of Xi-1 and Yi.
If we assume tropical profiles of N2O and its isotopocule ratios (e.g.,
δ15Nbulk) are determined purely by photochemistry with
initial mixing ratio of 320 ppb, a delta value of 0 ‰ and a
εsink of -50 ‰, then tropical air masses
vertically divided from Y8 through Y1 and X0 are expected to
line up on a solid line as portrayed in Fig. 7. Next, let us consider that
air mass X0 is mixed with Y1 to form X1. Based on the mass
balance of isotopocules before and after mixing, the resulting composition of
X1 is obtained as a curve, as shown in red in Fig. 7. Assuming
arbitrarily that the mixing ratio of Y1 to X1 is 0.1, and repeating
such mixing stepwise, then we obtain mixing ratio and isotope ratios of
N2O in X1 through X8 as black stars in Fig. 6. This
hypothetical, continuous mixing produces a curve that is qualitatively
consistent with observations made over the mid-latitudes or high latitudes.
The mixing effect must also be the cause of smaller ε in the
equatorial lower stratosphere (Fig. 4). The mean age of air deduced from
CO2 mixing ratio is known to be significantly larger than the phase lag
of the water vapor mixing ratio, a so-called tape recorder signal, in the
tropical stratosphere, which is explainable by mixing of old air from the
extratropics into the tropics (Waugh and Hall, 2002). In addition, the
difference in age between the Equator and mid-latitude (over Japan) decreases
concomitantly with decreasing altitude (Sugawara et al., unpublished data),
suggesting that the timescale of meridional mixing or transport is smaller
in the lower stratosphere than in the middle stratosphere, as suggested by
results of an earlier study (Boering et al., 1996).
Comparison with chemical transport model (ACTM) simulation
Presentation of Xi (black stars) obtained using the mixing
model with assumed Yi in the Rayleigh plot. The straight line shows
tropical vertical isotopocule fractionation without transport/mixing effect.
Curves show mixing between Xi-1 and Yi, where mixing ratio
Yi/Xi is assumed to be 0.1.
We further examined the importance of transport using an ACTM. Figure 8
presents results of the ACTM simulation with observational data. Although the
model approximates the photolysis of N2O in the longer wavelength region
(λ > 200 nm) with lower spectral resolution, profiles of the
N2O mixing ratio and isotopocule ratios were reproduced well, except in
the winter polar stratosphere, where dynamic processes specific to the polar
vortex might not be simulated appropriately in the model. In Fig. 9, the
model simulation and observations are compared on a Rayleigh plot. Again, the
model reproduced the difference between tropical and mid-latitudes or high
latitudes. Because in situ εphotolysis used in the
model calculation is nearly the same between low and high latitudes
(Fig. S6c), this agreement supports the inference that the major causes of
the difference are transport and mixing, which was previously suggested by
observations in the high latitudes (Park et al., 2004) and by 1-D or 3-D
model studies (McLinden et al., 2003; Morgan et al., 2004).
Comparison of vertical profiles of mixing ratio (a) and
δ15Nbulk (b) of N2O between observations
and simulation by the ACTM. Model simulations for equatorial profiles were
conducted for two dates because the observations were conducted during a 5- or
7-day period.
Comparison of results of ACTM simulation and stratospheric
observation in Rayleigh plot. The square region shown by broken lines in
panel (a) is enlarged in panel (b).
Vertical profiles of ratio of ε values for
15Nbulk and 18O (ψ) and the ratio of ε values for 15Nα and 15Nβ (η)
calculated in the manner similar to that of Kaiser et al. (2006). Grey bands
show values obtained by laboratory broadband photolysis experiments (Kaiser
et al., 2002b, 2003) and photooxidation experiments (Kaiser et al., 2002a;
Toyoda et al., 2004) with widths representing their uncertainty. For
photolysis, uncertainties associated with temperature (190–240 K) and
wavelength (190–220 nm) dependency are shown with dark and light grey,
respectively.
Share of photolysis and photooxidation
Kaiser et al. (2006) used the ratio of ε values for
15Nbulk and 18O (ψ) and the ratio of ε values for 15Nα and 15Nβ (η) to
estimate the relative share of photolysis and photooxidation based on the
fact that ψ and η are almost independent of transport processes
and are significantly different between the two decomposition processes. They
computed ψ and η values directly for each individual sample in
order to avoid statistical errors associated with linear regression to the
δ–δ plot which was adopted by Toyoda et al. (2004) and Park
et al. (2004). In Fig. 10, we show ψ and η values calculated
using the data presented in Fig. 2 in the manner similar to that of Kaiser et
al. (2006) except that we used the individual date of stratospheric entry for
each data point to normalize the δ values instead of using a single
tropopause date. Although it is noteworthy that errors in ψ and η
values increase concomitantly with decreasing altitude because of the
decrease in the δ values, low values are obtained just above the TTL
over the Equator (EQP, z=20 km) just as they are at other latitudes.
This result is in accordance with the indication by Kaiser et al. (2006) that
the photooxidation sink has a much larger fraction than 10 % in the lower
stratosphere. Although the loss rate of N2O in the lower stratosphere is
very slow and the majority of N2O injected into the stratosphere is
photolyzed in the middle stratosphere as noted by Park et al. (2004), the
share of photooxidation in in situ total loss increases in the lower
stratosphere (Fig. S6b). Therefore, there is a possibility of additional
decomposition of remaining N2O during the transport from the tropics to
the extratropics. In the transport which must be slower than the rate of
tropical upwelling, the isotopic signature of O(1D) pathway could be
imprinted, and the photochemically aged air mass could be transported into
the lower stratosphere of the tropics and extratropics. However, Morgan et
al. (2003) reported that inclusion of isotope fractionation for
photooxidation into their 2-D model does not make a significant contribution
to overall fractionation in the stratosphere, and Park et al. (2004)
discussed an alternative modeling approach with and without O(1D) sink
to test the importance of O(1D) reaction. Further studies using 3-D
model would be necessary to solve this controversial problem.