Nitrogen isotope fractionations between nitrogen oxides (NO and NO2) play a significant role in determining the nitrogen isotopic compositions (δ15N) of atmospheric reactive nitrogen. Both the equilibrium isotopic exchange between NO and NO2 molecules and the
isotope effects occurring during the NOx photochemical cycle are
important, but both are not well constrained. The nighttime and daytime
isotopic fractionations between NO and NO2 in an atmospheric simulation
chamber at atmospherically relevant NOx levels were measured. Then, the
impact of NOx level and NO2 photolysis rate on the combined
isotopic fractionation (equilibrium isotopic exchange and photochemical
cycle) between NO and NO2 was calculated. It was found that the
isotope effects occurring during the NOx photochemical cycle can be
described using a single fractionation factor, designated the Leighton cycle
isotope effect (LCIE). The results showed that at room temperature, the
fractionation factor of nitrogen isotopic exchange is 1.0289±0.0019,
and the fractionation factor of LCIE (when O3 solely controls the
oxidation from NO to NO2) is 0.990±0.005. The measured LCIE
factor showed good agreement with previous field measurements, suggesting
that it could be applied in an ambient environment, although future work is
needed to assess the isotopic fractionation factors of NO+RO2/HO2→NO2. The results were used to model the
NO–NO2 isotopic fractionations under several NOx conditions. The
model suggested that isotopic exchange was the dominant factor when NOx>20 nmol mol-1, while LCIE was more important at low
NOx concentrations (<1 nmol mol-1) and high rates of
NO2 photolysis. These findings provided a useful tool to quantify the
isotopic fractionations between tropospheric NO and NO2, which can be
applied in future field observations and atmospheric chemistry models.
Introduction
The nitrogen isotopic composition (δ15N) of reactive nitrogen
compounds in the atmosphere is an important tool in understanding the
sources and chemistry of atmospheric NOx (NO+NO2). It has been
suggested that the δ15N value of atmospheric nitrate
(HNO3, nitrate aerosols and nitrate ions in precipitation and snow)
imprints the δ15N value of NOx sources (Elliott et al.,
2009; Kendall et al., 2007); thus many studies have used the
δ15N values of atmospheric nitrate to investigate NOx sources
(Chang et al., 2018; Felix et al., 2012; Felix and Elliott, 2014; Gobel et
al., 2013; Hastings et al., 2004, 2009; Morin et al., 2009; Park et al.,
2018; Walters et al., 2015, 2018). However, there remain questions about how
isotopic fractionations that may occur during photochemical cycling of
NOx could alter the δ15N values as it partitions into
NOy (NOy= atmospheric nitrate, NO3, N2O5, HONO,
etc.; Chang et al., 2018; Freyer, 1991; Hastings et al., 2004; Jarvis et
al., 2008; Michalski et al., 2005; Morin et al., 2009; Zong et al., 2017).
Similarly, other complex reactive nitrogen chemistry, such as nitrate
photolysis and redeposition in ice and snow (Frey et al., 2009), may impact
the δ15N of NOy and atmospheric nitrate. The fractionation
between NO and NO2 via isotope exchange has been suggested to be the
dominant factor in determining the δ15N of NO2 and
ultimately atmospheric nitrate (Freyer, 1991; Freyer et al., 1993; Savarino
et al., 2013; Walters et al., 2016). However, isotopic fractionations occur
in most, if not all, NOx and NOy reactions, while most of these
are still unknown or, if calculated (Walters and Michalski, 2015),
unverified by experiments. Since the atmospheric chemistry of NOy
varies significantly in different environments (e.g., polluted vs. pristine,
night vs. day), the isotopic fractionations associated with NOy
chemistry are also likely to vary in different environments. These unknowns
could potentially bias conclusions about NOx source apportionment
reached when using nitrogen isotopes. Therefore, understanding the isotopic
fractionations between NO and NO2 during photochemical cycling could
improve our understanding of the relative role of sources versus chemistry
for controlling the δ15N variations in atmospheric NO2 and
nitrate.
In general, there are three types of isotopic fractionation effects
associated with NOx chemistry (Fig. 1a). The first type is the
equilibrium isotopic effect (EIE), i.e., isotope exchange between two
compounds without forming new molecules (Urey, 1947; Bigeleisen and Mayer,
1947), which for nitrogen isotopes in the NOx system is the 15NO+14NO2↔14NO+15NO2
exchange reaction (Begun and Melton, 1956; Walters et al., 2016). The second
type is the kinetic isotopic effect (KIE) associated with difference in
isotopologue rate coefficients during unidirectional reactions (Bigeleisen
and Wolfsberg, 1957). In the NOx system this KIE would manifest in the
oxidation of NO into NO2 by O3/HO2/RO2. The third type
is the photochemical isotope fractionation effect (PHIFE; Miller and Yung,
2000), which for NOx is the isotopic fractionation associated with
NO2 photolysis. All three fractionations could impact the δ15N value of NO2 and consequently atmospheric nitrate, but the
relative importance of each may vary.
(a) A sketch of the isotopic fractionation processes between NO and
NO2; both fractionation factors are determined in this work. (b) Results
from five dark experiments (red circles) yielded a line with slope of
28.1 ‰ and an α(NO2–NO) value of 1.0289,
while the results from five UV irradiation experiments (blue squares) showed
a smaller slope. (c) Results from five UV irradiation experiments (blue
squares) and a previous field study (purple triangle), comparing to the dark
experiments (red circle). The three lines represent different (α2-α1) values: the (α2-α1) =-10 ‰ line showed the lowest RMSE to our experimental
data as well as the previous field observations. The error bars in panels (b)
and (c) represented the combined uncertainties of NOx concentration
measurements and isotopic analysis.
The limited number of studies on the EIE in the NOx cycle have
significant uncertainties. Discrepancies in the EIE for 15NO+14NO2↔14NO+15NO2 have been
noted in several studies. Theoretical calculations predicted isotope
fractionation factors (α) ranging from 1.035 to 1.042 at room
temperature (Begun and Fletcher, 1960; Monse et al., 1969; Walters and
Michalski, 2015) due to the different approximations used to calculate
harmonic frequencies in each study. Likewise, two separate experiments
measured different room temperature fractionation factors of 1.028±0.002 (Begun and Melton, 1956) and 1.0356±0.0015 (Walters et al.,
2016). A concern in both experiments is that they were conducted in small
chambers with high NOx concentrations (hundreds of micromoles per mole), significantly higher than typical ambient atmospheric NOx
levels (usually less than 0.1 µmol mol-1). Whether the isotopic
fractionation factors determined by these experiments are applicable in the
ambient environment is uncertain because of possible wall effects and
formation of higher oxides, notably N2O4 and N2O3 at
these high NOx concentrations.
Even less research has examined the KIE and PHIFE occurring during NOx
cycling. The KIE of NO+O3 has been theoretically calculated
(Walters and Michalski, 2016) but has not been experimentally verified. The
NO2 PHIFE has not been experimentally determined or theoretically
calculated. As a result, field observation studies often overlook the
effects of PHIFE and KIE. Freyer et al. (1993) measured NOx
concentrations and the δ15N values of NO2 over a 1-year
period at Julich, Germany, and inferred a combined NOx isotope
fractionation factor (EIE+KIE+PHIFE) of 1.018±0.001. Freyer et
al. (1993) suggested that the NOx photochemical cycle (KIE and PHIFE)
tends to diminish the equilibrium isotopic fractionation (EIE) between NO
and NO2. Even if this approach were valid, applying this single
fractionation factor elsewhere, where NOx and O3 concentrations and
actinic fluxes are different, would be tenuous given that these factors may
influence the relative importance of EIE, KIE and PHIFE (Hastings et al.,
2004; Walters et al., 2016). Therefore, to quantify the overall isotopic
fractionations between NO and NO2 at various tropospheric conditions,
it is crucial to know (1) isotopic fractionation factors of EIE, KIE and
PHIFE individually and (2) the relative importance of each factor under
various conditions.
In this work, we aim to quantify the nitrogen isotope fractionation factors
between NO and NO2 at photochemical equilibrium. First, we measure the
N isotope fractionations between NO and NO2 in an atmospheric
simulation chamber at atmospherically relevant NOx levels. Then, we
provide mathematical solutions to assess the impact of NOx level and
NO2 photolysis rate (j(NO2)) on the relative importance of EIE, KIE
and PHIFE. Subsequently we use the solutions and chamber measurements to
calculate the isotopic fractionation factors of EIE, KIE and PHIFE. Lastly,
using the calculated fractionation factors and the equations, we model the
NO–NO2 isotopic fractionations at several sites to illustrate the
behavior of δ15N values of NOx in the ambient environment.
Methods
The experiments were conducted using a 10 m3 atmospheric simulation
chamber at the National Center for Atmospheric Research (see descriptions in
Appendix A and Zhang et al., 2018). A set of mass flow controllers was used
to inject NO and O3 into the chamber. NO was injected at 1 L min-1
from an in-house NO/N2 cylinder (133.16 µmol mol-1 NO in
ultrapure N2), and O3 was generated by flowing zero air through a flow tube equipped with a UV Pen-Ray lamp (UVP LLC., CA)
into the chamber at 5 L min-1. NO and NO2 concentrations were monitored in real time
by chemiluminescence with a detection limit of 0.5 nmol mol-1 (model
CLD 88Y, Eco Physics, MI) as were O3 concentrations using a UV
absorption spectroscopy with a detection limit of 0.5 nmol mol-1 (model 49, Thermo Scientific, CO). In each experiment, the actual amounts of NO and
O3 injected were calculated using measured NOx and O3
concentrations after a steady state was reached (usually within 1 h). The wall
loss rate of NO2 was tested by monitoring O3 (29 nmol mol-1)
and NOx (62 nmol mol-1) over a 4 h period. After the NO and
NO2 concentrations reached a steady state, no decrease in NO2
concentrations was observed, showing that chamber wall loss was negligible.
Three experiments were conducted to measure the δ15N value of
the tank NO (i.e., the δ15N value of total NOx). In each
of these experiments, a certain amount of O3 was first injected into
the chamber, then approximately the same amount of NO was injected into the
chamber to ensure 100 % of the NOx was in the form of NO2 with
little O3 (<15 nmol mol-1) remaining in the chamber such
that the O3+NO2 reaction was negligible. The NO2 in the
chamber was then collected and its δ15N value measured, which
equates to the δ15N value of the tank NO.
Two sets of experiments were conducted to separately investigate the EIE,
KIE and PHIFE. The first set of experiments was conducted in the dark. In
each of these dark experiments, a range of NO and O3
([O3] < [NO]) was injected into the chamber to produce
NO–NO2 mixtures with [NO]/[NO2] ratios ranging from 0.43 to 1.17.
The N isotopes of these mixtures were used to investigate the EIE between NO
and NO2. The second set of experiments was conducted under irradiation
of UV lights (300–500 nm; see Appendix A for irradiation spectrum). Under
such conditions, NO, NO2 and O3 reached a photochemically steady
state, which combined the isotopic effects of EIE, KIE and PHIFE.
In all experiments, the concentrations of NO, NO2 and O3 were
allowed to reach a steady state, and the product NO2 was collected from
the chamber using a honeycomb denuder tube. After the NO, NO2 and
O3 concentrations reached a steady state, well-mixed chamber air was
drawn out through a 40 cm long Norprene thermoplastic tubing at 10 L min-1 and passed through a honeycomb denuder system (Chemcomb 3500,
Thermo Scientific). Based on flow rate, the NO2 residence time in the
tubing was less than 0.5 s; thus in the light-on experiments where NO and
O3 coexisted, the NO2 produced inside the transfer tube through
NO+O3 reactions should be <0.03 nmol mol-1 (using the
upper limit of NO and O3 concentrations in our experiments). The
honeycomb denuder system consisted of two honeycomb denuder tubes connected
in series. Each honeycomb denuder tube is a glass cylinder 38 mm long and 47 mm in diameter and consists of 212 hexagonal tubes with inner diameters of 2 mm. Before collecting samples, each denuder tube was coated with a solution
of 10 % KOH and 25 % guaiacol in methanol and then dried by flowing
N2 gas through the denuder tube for 15 s (Williams and Grosjean,
1990; Walters et al., 2016). The NO2 reacted with the guaiacol coating and
was converted into NO2- that was retained on the denuder tube wall
(Williams and Grosjean, 1990). NO was inert to the denuder tube coating: a
control experiment sampled pure NO using the denuder tubes, which did not
show any measurable NO2-. The NO2 collection efficiency of a
single honeycomb denuder tube was tested in another control experiment: air
containing 66 nmol mol-1 of NO2 was drawn out of the chamber
through a denuder tube, and the NO2 concentration at the exit of the
tube holder was measured and found to be below the detection limit
(<1 nmol mol-1), suggesting that the collection efficiency was
nearly 100 % when [NO2] <66 nmol mol-1. Furthermore,
when the denuder system consisted of two denuder tubes in series, NO2- in the second denuder was below the detection limit,
indicating trivial NO2 breakthrough. Each NO2 collection lasted
for 0.5–3 h in order to collect enough NO2- for isotopic
analysis (at least 300 nmol). After collection, the NO2- was
leached from each denuder tube by rinsing thoroughly with 10 mL deionized
water into a clean polypropylene container and stored frozen until isotopic
analysis. Isotopic analysis was conducted at the Purdue Stable Isotope
Laboratory. For each sample, approximately 50 nmol of the NO2-
extract was mixed with 2 M sodium azide solution in an acetic acid buffer in an
airtight glass vial, then shaken overnight to completely reduce all the
NO2- to N2O(g) (Casciotti and McIlvin, 2007; McIlvin
and Altabet, 2005). The product N2O was directed into a Thermo Scientific
GasBench equipped with a cryotrap, then the δ15N of the N2O
was measured using a Delta-V isotope ratio mass spectrometer (IRMS). Six coated
denuder tubes that did not get exposed to NO2 were also analyzed using
the same chemical procedure, which did not show any measurable signal on the
IRMS, suggesting that the blank from both the sampling process and the chemical
conversion process was negligible. The overall analytical uncertainty for
δ15N analysis was 0.5 ‰ (1σ) based
on replicate analysis of in-house NO2- standards.
Results and discussionsEquilibrium isotopic fractionation between NO and <inline-formula><mml:math id="M198" display="inline"><mml:mrow class="chem"><mml:msub><mml:mi mathvariant="normal">NO</mml:mi><mml:mn mathvariant="normal">2</mml:mn></mml:msub></mml:mrow></mml:math></inline-formula>
The equilibrium isotope fractionation factor, α(NO2–NO), is the
15N enrichment in NO2 relative to NO and is expressed as the
ratio of rate constants k2/k1 of two reactions:
R115NO2+14NO→15NO+14NO2,rate constant=k1R215NO+14NO2→15NO2+14NO,rate
constant=k2=k1α(NO2–NO),
where k1 is the rate constant of the isotopic exchange, which was
previously determined to be 8.14×10-14 cm3 s-1
(Sharma et al., 1970). The reaction time required for NO–NO2 to reach
isotopic equilibrium was estimated using the exchange rate constants in a
simple kinetics box model (BOXMOX; Knote et al., 2015). The model predicts
that at typical NOx concentrations used during the chamber experiments (7.7–62.4 nmol mol-1), isotopic equilibrium would be reached within 15 min (see Appendix B). Since the sample collection usually started 1 h after NOx was well mixed in the chamber, there was sufficient time to reach full isotope equilibrium. The isotope equilibrium fractionation factor
(α(NO2–NO)) is then calculated to be
α(NO2–NO)=[15NO2][14NO][14NO2][15NO]=R(NO2)R(NO),
where R(NO,NO2) are the 15N/14N ratios of NO and
NO2. By definition, the δ15N(NO) is (R(NO)/R(reference)-1)×1000 ‰, and δ15N(NO2) is (R(NO2)/R(reference)-1)×1000 ‰, but hereafter, the δ15N values of NO,
NO2 and NOx are referred to as δ(NO), δ(NO2)
and δ(NOx), respectively. Equation (1) leads to
δNO2-δNO=αNO2–NO-1(1+δ(NO)).
Using Eq. (2) and applying NOx isotopic mass balance (δ(NOx)=f(NO2)δ(NO2)+(1-f(NO2))δ(NO), f(NO2)=[NO2]/([NO]+[NO2])) yields
δNO2-δNOx1+NO2=αNO2–NO-1αNO2–NO1-fNO2.
Here, δ(NOx) equals the δ15N value of the
cylinder NO, and f(NO2) is the molar fraction of NO2 with respect
to total NOx. Three experiments (Table 1) that measured δ(NOx) showed consistent δ(NOx) values of -58.7±0.8 ‰ (n=3), indicating that δ(NOx)
remained unchanged throughout the experiments (as expected for isotope mass
balance). Thus, the δ(NOx) can be treated as a constant in Eq. (3), and the linear regression of (δ(NO2)-δ(NOx))/(1+δ(NO2)) versus 1-f(NO2) should have an
intercept of 0 and a slope of (α(NO2–NO)-1)/α(NO2–NO).
Experimental conditions; concentrations of NO, NO2 and O3 at a steady state; and measured δ(NO2) values.
The plot of (δ(NO2)-δ(NOx))/(1+δ(NO2)) as a function of 1-f(NO2) values from five experiments
yields an α(NO2–NO) value of 1.0289±0.0019 at room
temperature (Fig. 1b and Table 1). This fractionation factor is comparable
to previously measured values but with some differences. Our result agrees
well with the α(NO2–NO) value of 1.028±0.002
obtained by Begun and Melton (1956) at room temperature. However, Walters et
al. (2016) determined the α(NO2–NO) values of NO–NO2
exchange in a 1 L reaction vessel, which showed a slightly higher
α(NO2–NO) value of 1.035. This discrepancy might originate from
rapid heterogeneous reactions on the wall of the reaction vessel at high
NOx concentrations and the small chamber size used by Walters et al. (2016). They used a reaction vessel made of Pyrex, which is known to absorb
water (Do Remus et al., 1983; Takei et al., 1997) and can react with
NO2, forming HONO, HNO3 and other N compounds. Additionally,
previous studies have suggested that Pyrex walls enhance the formation rate
of N2O4 by over an order of magnitude (Barney and
Finlayson-Pitts, 2000; Saliba et al., 2001), which at isotopic equilibrium
is enriched in 15N compared to NO and NO2 (Walters and Michalski,
2015). Therefore, their measured α(NO2–NO) might be slightly
higher than the actual α(NO2–NO) value. In this work, the 10 m3 chamber has a much smaller surface-to-volume ratio relative to
Walters et al. (2016), which minimizes wall effects, and the walls were made
of Teflon, which minimizes NO2 surface reactivity, as evidenced by the NO2 wall loss control experiment. Furthermore, the low NOx
mixing ratios in our experiments minimized N2O4 and N2O3
formation. At NO and NO2 concentrations of 50 nmol mol-1, the
steady-state concentrations of N2O4 and N2O3 were
calculated to be 0.014 and 0.001 pmol mol-1, respectively (Atkinson et
al., 2004). Therefore, we suggest that our measured α(NO2–NO) value
(1.0289±0.0019) may better reflect the room temperature (298 K)
NO–NO2 EIE in the ambient environment.
Unfortunately, the chamber temperature could not be controlled, so we were
not able to investigate the temperature dependence of the EIE. Hence, we
speculate that the α(NO2–NO) follows a similar temperature
dependence pattern calculated in Walters et al. (2016). Walters et al. (2016) suggested that the α(NO2–NO) value would be 0.0047
higher at 273 K and 0.002 lower at 310 K relative to room temperature (298 K). Using this pattern and our experimentally determined data, we suggest that the α(NO2–NO) values at 273, 298 and 310 K are
1.0336±0.0019, 1.0289±0.0019 and 1.0269±0.0019,
respectively. This 0.0067 variation at least partially contributes to the
daily and seasonal variations in δ15N values of NO2 and
nitrate in some areas (e.g., polar regions with strong seasonal temperature
variation). Thus, future investigations should be conducted to verify the
EIE temperature dependence.
Kinetic isotopic fractionation of Leighton cycle
The photochemical reactions of NOx will compete with the isotope
exchange fractionations between NO and NO2. The NO–NO2
photochemical cycle in the chamber was controlled by the Leighton cycle:
NO2 photolysis and the NO+O3 reaction. This is because there
were no volatile organic compounds (VOCs) in the chamber, so no RO2 was produced, which excludes the
NO+RO2 reaction. Likewise, the low water vapor content (relative humidity <10 %) and the minor flux of photons (<310 nm) results in minimal OH
production and hence little HO2 formation, and subsequently a trivial
amount of NO2 would be formed by NO+HO2. Applying these
limiting assumptions, the EIE between NO and NO2 (Reactions R1–R2) is only
competing with the KIE (Reactions R3–R4) and the PHIFE in Reactions (R5)–(R6):
R314NO2→14NO+O,rate constant=j(NO2)R415NO2→15NO+O,rate constant=j(NO2)α1R514NO+O3→14NO2+O2,rate constant=k5R615NO+O3→15NO2+O2,rate constant=k5α2,
in which j(NO2) is the NO2 photolysis rate (1.4×10-3 s-1 in these experiments), k5 is the rate constant for the
NO+O3 reaction (1.73×10-14 cm3 s-1; Atkinson
et al., 2004), and α1,2 are isotopic fractionation factors for
the two reactions. Previous studies (Freyer et al., 1993; Walters et al.,
2016) have attempted to assess the competition between EIE (Reactions R1–R2), KIE and
PHIFE (Reactions R3–R6), but none of them quantified the relative importance of the
two processes, nor were α1 or α2 values
experimentally determined. Here we provide the mathematical solution of EIE,
KIE and PHIFE to illustrate how Reactions (R1)–(R6) affect the isotopic fractionations
between NO and NO2.
First, the NO2 lifetime with respect to isotopic exchange with NO
(τexchange) and photolysis (τphoto) was determined:
4τexchange=1k1[NO]5τphoto=1j(NO2).
We then define an A factor:
A=τexchangeτphotowhen j(NO2)≠00when j(NO2)=0.
Using Reactions (R1)–(R6) and Eqs. (1)–(6), we solved steady-state δ(NO2) and
δ(NO) values (see calculations in Appendix C). Our calculations show
that the δ(NO2)–δ(NO) and δ(NO2)–δ(NOx) values at a steady state can be expressed as functions of α1, α2, α(NO2–NO) and A:
7δNO2-δNO=α2-α1A+αNO2–NO-1α2A+αNO2–NO1+δNO2≈α2-α1A+αNO2–NO-1A+11+δNO28δNO2-δNOx=α2-α1A+αNO2–NO-1α2A+αNO2–NO1+δNO21-fNO2≈α2-α1A+αNO2–NO-1A+11+δNO21-fNO2.
Equation (7) shows the isotopic fractionation between NO and NO2;
δ(NO2)–δ(NO) is mainly determined by A; the EIE
factor (α(NO2–NO)-1) and the (α2-α1)
factor assuming 1+δ(NO2) is close to 1. This (α2-α1) represents a combination of KIE and PHIFE,
suggesting that they act together as one factor; therefore, we name the (α2-α1) factor the Leighton cycle isotopic effect (LCIE).
Using measured δ(NO2) and δ(NOx) values, A values
(Table 1), and the previously determined α(NO2–NO) value, we plot δNO2-δNOx1+δNO21-f(NO2)equals δNO2-δNO1+δNO2 against the A value and use
Eqs. (7) and (8) to estimate the (α2-α1)
value (Fig. 1c). The plot shows that the best fit for the LCIE factor is
(-10±5) ‰ (root mean square error, RMSE, was
lowest when α2-α1=-10 ‰).
The uncertainties in the LCIE factor are relatively higher than that of the
EIE factor, mainly because of the accumulated analytical uncertainties at
low NOx and O3 concentrations and low A values (0.10–0.28) due to
the relatively low j(NO2) value (1.4×10-3 s-1) under the chamber irradiation conditions.
This LCIE factor determined in our experiments is in good agreement with
theoretical calculations. Walters and Michalski (2016) previously used an
ab initio approach to determine an α2 value of 0.9933 at room
temperature, 0.9943 at 237 K and 0.9929 at 310 K. The total variation in
α2 values from 273 to 310 K is only 1.4 ‰, significantly smaller than our experimental uncertainty (±5 ‰). The α1 value was calculated using a
zero-point energy (ZPE) shift model (Miller and Yung, 2000) to calculate the isotopic
fractionation of NO2 by photolysis. Briefly, this model assumes both
isotopologues have the same quantum yield function, and the PHIFE was only
caused by the differences in the 15NO2 and 14NO2
absorption cross section as a function of wavelength; thus α1
values do not vary by temperature. The 15NO2 absorption
cross section was calculated by shifting the 14NO2 absorption
cross section by the 15NO2 zero-point energy (Michalski et al.,
2004). When the ZPE shift model was used with the irradiation spectrum of
the chamber lights, the resulting α1 value was 1.0023.
Therefore, the theoretically predicted α2-α1 value
should be -0.0090, i.e., -9.0±0.7 ‰ when
the temperature ranges from 273 to 310 K. This result shows excellent
agreement with our experimentally determined room temperature α2-α1 value of -10±5 ‰ .
This model was then used to evaluate the variations in α1 under different lighting conditions. The tropospheric ultraviolet and visible (TUV) radiation model (TUV5.3.2; Madronich and
Flocke, 1999) was used to calculate the solar wavelength spectrum at three
different conditions: early morning or late afternoon (solar zenith angle =85∘), midmorning or afternoon (solar zenith angle =45∘), and 12:00 local time (LT; solar
zenith angle =0∘). These spectra were used in the ZPE shift model
to calculate the α1 values, which are 1.0025, 1.0028 and
1.0029 at solar zenith angles of 85, 45 and 0∘, respectively. These
values, along with the predicted α1 value in the chamber,
showed a total span of 0.6 ‰ (1.0026±0.0003),
which is again significantly smaller than our measured uncertainty.
Therefore, we suggest that our experimentally determined LCIE factor
(-10±5 ‰) can be used in most tropospheric
solar irradiation spectra.
The equations can also be applied in tropospheric environments to calculate
the combined isotopic fractionations of EIE and LCIE for NO and NO2.
First, the NO2 sink reactions (mainly NO2+OH in the daytime) are
at least 2–3 orders of magnitude slower than the Leighton cycle and the
NO–NO2 isotope exchange reactions (Walters et al., 2016); therefore
their effects on the δ(NO2) should be minor. Second, although
the conversion of NO into NO2 in the ambient environment is also
controlled by NO+RO2 and HO2 in addition to NO+O3
(e.g., King et al., 2001), Eq. (7) still showed good agreement with field
observations in previous studies. Freyer et al. (1993) determined the annual
average daytime δ(NO2)–δ(NO) at Julich, Germany, along
with average daytime NO concentration (9 nmol mol-1, similar to our
experimental conditions) to be +18.03±0.98 ‰. Using Eq. (7), assuming the daytime average j(NO2) value throughout
the year was (5.0±1.0)×10-3, and a calculated A value
from measured NOx concentration ranged from 0.22 to 0.33, the average
NO–NO2 fractionation factor was calculated to be +19.8±1.4 ‰ (Fig. 1c), in excellent agreement with the
measurements in the present study. This agreement suggests that the
NO+RO2/HO2 reactions might have similar fractionation factors as
NO+O3. Therefore, we suggest that Eqs. (7) and (8) can be used to estimate
the isotopic fractionations between NO and NO2 in the troposphere.
First, Eq. (7) was used to calculate the Δ(NO2–NO)=δ(NO2)–δ(NO) at a wide range of NOx concentrations and f(NO2) and j(NO2) values (Fig. 2a–d), assuming 1+δ(NO2)≈1. j(NO2) values of 0 s-1 (Fig. 2a),
1.4×10-3 s-1 (Fig. 2b), 5×10-3 s-1
(Fig. 2c) and 1×10-2 s-1 (Fig. 2d) were selected to
represent nighttime, dawn (as well as the laboratory conditions of our
experiments), daytime average and 12:00 LT, respectively. Each panel represented
a fixed j(NO2) value, and the Δ(NO2–NO) values were calculated as a function of the A value, which was derived from NOx concentration and f(NO2). The A values have a large span, from 0 to 500, depending on the j(NO2) value and the NO concentration. When A=0
(j(NO2)=0) and f(NO2)<1 (meaning NO and NO2 coexist and [O3] = 0), Eqs. (7) and (8) become Eqs. (2) and (3), showing that the EIE was the sole factor; the Δ(NO2–NO) values were solely controlled by
EIE, which has a constant value of +28.9 ‰ at 298 K
(Fig. 2a). When j(NO2)>0, the calculated Δ(NO2–NO) values showed a wide range from -10.0 ‰
(controlled by LCIE factor: α2-α1=-10 ‰) to +28.9 ‰ (controlled by EIE
factor: α(NO2–NO)-1=+28.9 ‰). Figure 2b–d display the transition from an LCIE-dominated regime to an EIE-dominated
regime. The LCIE-dominated regime is characterized by low [NOx]
(<50 pmol mol-1), representing remote ocean areas and polar
regions (Beine et al., 2002; Custard et al., 2015). At this range the A value can be greater than 200; thus Eq. (7) can be simplified as Δ(NO2–NO)=α2-α1, suggesting that the LCIE
almost exclusively controls the NO–NO2 isotopic fractionation. The
Δ(NO2–NO) values of these regions are predicted to be <0 ‰ during most times of the day and <-5 ‰ at 12:00 LT. On the other hand, the EIE-dominated regime
was characterized by high [NOx] (>20 nmol mol-1) and
low f(NO2) (<0.6), representative of regions with intensive
NO emissions, e.g., near roadside or stack plumes (Clapp and Jenkin, 2001;
Kimbrough et al., 2017). In this case, the τexchange are
relatively short (10–50 s) compared to the τphoto (approximately
100 s at 12:00 LT and 1000 s at dawn); therefore the A values are small
(0.01–0.5). The EIE factor in this regime thus is much more important than
the LCIE factor, resulting in high Δ(NO2–NO) values
(>20 ‰). Between the two regimes, both EIE
and LCIE are competitive, and therefore it is necessary to use Eq. (7) to
quantify the Δ(NO2–NO) values.
Calculating isotopic fractionation values between NO–NO2
(Δ(NO2–NO); a–d) and NOx–NO2 (Δ(NO2–NOx); e–h) at various j(NO2), NOx level and
f(NO2) using Eqs. (7) and (8). Each panel represents a fixed
j(NO2) value (on the upper-right side of each panel), and the
fractionation values are shown by color. Lines are contours with the same
fractionation values at an interval of 5 ‰; the contour
line representing 0 ‰ was marked in each panel except for
panels (a) and (e).
Figure 2 also implies that changes in the j(NO2) value can cause the
diurnal variations in Δ(NO2–NO) values. Changing j(NO2)
would affect the value of A and consequently the NO–NO2 isotopic
fractionations in two ways: (1) changes in the j(NO2) value would change the
photolysis intensity and therefore the τphoto value; (2) in addition,
changes in the j(NO2) value would also alter the steady-state NO
concentration, therefore changing the τexchange (Fig. 2c). The
combined effect of these two factors on the A value varies along with the
atmospheric conditions and thus needs to be carefully calculated using
NOx concentration data and atmospheric chemistry models.
We then calculated the differences in δ15N values between
NO2 and total NOx, e.g., Δ(NO2–NOx)=δ(NO2)–δ(NOx) in Fig. 2e–h. Since Δ(NO2–NOx) are connected through the observed δ15N of
NO2 (or nitrate) to the δ15N of NOx sources, this
term might be useful in field studies (e.g., Chang et al., 2018; Zong et
al., 2017). The calculated Δ(NO2–NOx) values (Fig. 2e–h)
also showed an LCIE-dominated regime at low [NOx] and an EIE-dominated
regime at high [NOx]. The Δ(NO2–NOx) values were
dampened by the 1-f(NO2) factor comparing to Δ(NO2–NO),
as shown in Eqs. (3) and (8): Δ(NO2–NOx)=Δ(NO2–NO) (1-f(NO2)). At high f(NO2) values (>0.8), the differences between δ(NO2) and δ(NOx)
were less than 5 ‰; thus the measured δ(NO2) values were similar to δ(NOx), although the
isotopic fractionation between NO and NO2 could be noteworthy. Some
ambient environments with significant NO emissions or high NO2
photolysis rates usually have f(NO2) values between 0.4 and 0.8 (Mazzeo
et al., 2005; Vicars et al., 2013). In this scenario, the Δ(NO2–NOx) values in Fig. 2f–h showed wide ranges of -4.8 ‰ to +15.6 ‰, -6.0 ‰ to +15.0 ‰ and -6.3 ‰ to +14.2 ‰ at
j(NO2)=1.4×10-3, 5×10-3 and 1×10-2 s-1, respectively. These significant
differences again highlighted the importance of both LCIE and EIE (Eqs. 7
and 8) in calculating the Δ(NO2–NOx). In the following
discussion, we assume that (1) the α1 value remains constant (see
discussion above), (2) the NO+RO2/HO2 reactions have the same
fractionation factors (α2) as NO+O3, and (3) both EIE and
LCIE do not display significant temperature dependence. We then use Eqs. (7) and (8) and this laboratory-determined LCIE factor (-10 ‰) to calculate the nitrogen isotopic fractionation
between NO and NO2 at various tropospheric atmospheric conditions.
Implications
The daily variations in Δ(NO2–NOx) values at two roadside
NOx monitoring sites were predicted to demonstrate the effects of
NOx concentrations to the NO–NO2 isotopic fractionations. Hourly
NO and NO2 concentrations were acquired from a roadside site at
Anaheim, CA (https://www.arb.ca.gov, last access: 9 August 2019), and an urban site at Evansville, IN
(http://idem.tx.sutron.com, last access: 9 August 2019), on 25 July 2018. The hourly j(NO2) values
output from the TUV model (Madronich and Flocke, 1999) at these locations
were used to calculate the daily variations in Δ(NO2–NOx)
values (Fig. 3a, b) by applying Eq. (8) and assuming 1+δ(NO2)≈1. Hourly NOx concentrations were 12–51 nmol mol-1 at Anaheim and 9–38 nmol mol-1 at Evansville, and the
f(NO2) values at both sites did not show significant daily variations
(0.45±0.07 at Anaheim and 0.65±0.08 at Evansville), likely
because the NOx concentrations were controlled by the high NO emissions
from the road (Gao, 2007). The calculated Δ(NO2–NOx)
values using Eq. (8) showed significant diurnal variations. During the
nighttime, the isotopic fractionations were solely controlled by the EIE; the predicted Δ(NO2–NOx) values were +14.5±2.0 ‰ and +8.7±2.1 ‰ at
Anaheim and Evansville, respectively. During the daytime, the existence of
LCIE lowered the predicted Δ(NO2–NOx) values to
+9.8±1.7 ‰ at Anaheim and +3.1±1.5 ‰ at Evansville, while the f(NO2) values at both
sites remained similar. The lowest Δ(NO2–NOx) values for
both sites (+7.0 ‰ and +1.7 ‰)
occurred around 12:00 LT, when the NOx photolysis was the most intense. In
contrast, if one neglects the LCIE factor in the daytime, the Δ(NO2–NOx) values would be +12.9±1.5 ‰ and +10.0±1.6 ‰, respectively, an overestimation of 3.1 ‰ and 6.9 ‰. These discrepancies suggested that the LCIE played
an important role in the NO–NO2 isotopic fractionations, and neglecting
it could bias the NOx source apportionment using δ15N of
NO2 or nitrate.
NOx concentrations and calculated Δ(NO2–NOx)
values at four sites. Stacked bars show the NO and NO2 concentrations
extracted from monitoring sites (a–c) or calculated using a 0-D box model (d);
the red lines are Δ(NO2–NOx) values at each site. Note
that the NOx concentration (left y axis in panel d) is different from
the rest.
The role of LCIE was more important in less polluted sites. The Δ(NO2–NOx) values were calculated for a suburban site near San Diego,
CA, USA, again using the hourly NOx concentrations
(https://www.arb.ca.gov; Fig. 3c) and j(NO2) values calculated from the
TUV model. NOx concentrations at this site varied from 1 to 9 nmol mol-1, assuming 1+δ(NO2)≈1. During the
nighttime, NOx was in the form of NO2 (f(NO2=1)
because O3 concentrations were higher than NOx; thus the δ(NO2) values should be identical to δ(NOx) (Δ(NO2–NOx)=0). In the daytime a certain amount of NO was
produced by direct NO emission and NO2 photolysis, but the f(NO2)
was still high (0.73±0.08). Our calculation suggested that the daytime
Δ(NO2–NOx) values should be only +1.3±3.2 ‰, with a lowest value of -1.3 ‰.
These Δ(NO2–NOx) values were similar to the observed and
modeled summer daytime δ(NO2) values in West Lafayette, IN
(Walters et al., 2018), which suggest that the average daytime Δ(NO2–NOx) values at NOx=3.9±1.2 nmol mol-1 should range from +0.1 ‰ to +2.4 ‰. In this regime, we suggest that the Δ(NO2–NOx) values were generally small due to the significant
contribution of LCIE and high f(NO2).
The LCIE should be the dominant factor controlling the NO–NO2 isotopic
fractionation in remote regions, resulting in a completely different diurnal
pattern of Δ(NO2–NOx) compared with the urban–suburban
area. Direct hourly measurements of NOx at remote sites are rare; thus
we used a total NOx concentration of 50 pmol mol-1 and a daily O3
concentration of 20 nmol mol-1 at Summit, Greenland (Dibb et al., 2002;
Hastings et al., 2004; Honrath et al., 1999; Yang et al., 2002) and assumed 1+δ(NO2)≈1 and that the conversion of NO to NO2
was completely controlled by O3 to calculate the NO/NO2 ratios.
Here the isotopes of NOx were almost exclusively controlled by the LCIE
due to the high A values (>110). The Δ(NO2–NOx) values displayed a clear diurnal pattern (Fig. 3d), with
a maximum value of -0.3 ‰ in the “nighttime” (solar
zenith angle >85∘) and a minimum value of -5.0 ‰ during midday. This suggests that the isotopic
fractionations between NO and NO2 were almost completely controlled by
LCIE in remote regions when NOx concentrations were <0.1 nmol mol-1. However, since the isotopic fractionation factors of
nitrate formation reactions (NO2+OH, NO3+HC,
N2O5+H2O) are still unknown, more studies are needed to
fully explain the daily and seasonal variations in δ(NO3-)
in remote regions.
Nevertheless, our results have a few limitations. First, currently there are
very few field observations that can be used to evaluate our model;
therefore, future field observations that measure the δ15N
values of ambient NO and NO2 should be carried out to test our model.
Second, more work, including theoretical and experimental studies, is needed
to investigate the isotope fractionation factors occurring during the
conversion from NOx to NOy and nitrate: in the NOy cycle, EIE
(isotopic exchange between NO2, NO3 and N2O5), KIE
(formation of NO3, N2O5 and nitrate) and PHIFE (photolysis of
NO3, N2O5, HONO and sometimes nitrate) may also exist and be
relevant for the δ15N of HNO3 and HONO. In particular, the
N isotope fractionation occurring during the NO2+OH→HNO3 reaction needs investigation. Such studies could help us to model the isotopic fractionation between NOx emission and nitrate and
eventually enable us to analyze the δ15N value of NOx
emission by measuring the δ15N values of nitrate aerosols and
nitrate in wet depositions. Third, our discussion only focuses on the
reactive nitrogen chemistry in the troposphere; however, the nitrogen
chemistry in the stratosphere is drastically different from the tropospheric
chemistry; thus future studies are also needed to investigate the isotopic
fractionations in the stratospheric nitrogen chemistry. Last, the
temperature dependence of both EIE and LCIE needs to be carefully
investigated because of the wide range of temperature in both the troposphere
and stratosphere. Changes in temperature could alter the isotopic
fractionation factors of both EIE and LCIE as well as contribute to the
seasonality of isotopic fractionations between NOx and NOy
molecules.
Conclusions
The effect of NOx photochemistry on the nitrogen isotopic
fractionations between NO and NO2 was investigated. We first measured
the isotopic fractionations between NO and NO2 and provided
mathematical solutions to assess the impact of NOx level and NO2
photolysis rate (j(NO2)) on the relative importance of EIE and LCIE. The
EIE and LCIE isotope fractionation factors at room temperature were
determined to be 1.0289±0.0019 and 0.990±0.005, respectively.
These calculations and measurements can be used to determine the steady-state Δ(NO2–NO) and Δ(NO2–NOx) values at room
temperature. Subsequently we applied our equations to polluted, clean and
remote sites to model the daily variations in Δ(NO2–NOx)
values. We found that the Δ(NO2–NOx) values could vary
from over +20 ‰ to less than -5 ‰
depending on the environment: in general, the role of LCIE becomes more
important at low NOx concentrations, which tend to decrease the Δ(NO2–NOx) values. Our work provided a mathematical approach to
quantify the nitrogen isotopic fractionations between NO and NO2 that
can be applied to many tropospheric environments, which could help interpret
the measured δ15N values of NO2 and nitrate in field
observation studies.
Chamber descriptions
The chamber is a 10 m3 Teflon bag equipped with several standard
instruments including a temperature and humidity probe, NOx monitor and
O3 monitor. A total of 128 wall-mounted blacklight tubes surrounded the chamber to
mimic tropospheric photochemistry, and the photolysis rate of NO2
(j(NO2)) when all lights are on has been previously determined to be
1.4×10-3 s-1, similar to a j(NO2) coefficient at an
81∘ solar zenith angle. The irradiation spectrum of the black lights
is shown in Fig. A1. The chamber was kept at room temperature and 1 atm. Before each experiment, the chamber was flushed with
zero air at 40 L min-1 for at least 12 h to ensure the background
NOx, O3 and other trace gases were below the detection limit.
Spectral actinic flux versus wavelengths of the UV light source
used in our experiments.
Box model assessing the time needed for NO–<inline-formula><mml:math id="M827" display="inline"><mml:mrow class="chem"><mml:msub><mml:mi mathvariant="normal">NO</mml:mi><mml:mn mathvariant="normal">2</mml:mn></mml:msub></mml:mrow></mml:math></inline-formula> to reach
isotopic equilibrium
The time needed to reach NO–NO2 isotopic equilibrium during light-off
experiments was assessed using a 0-D box model. This box model contains
only two reactions:
BR115NO2+14NO→15NO+14NO2k=8.14000×10-14cm3s-1BR215NO+14NO2→15NO2+14NOk′=8.37525×10-14cm3s-1,
where k and k′ are rate constants of the reactions. The differences in rate
constants were calculated by assuming an α(NO2–NO) value of
1.0289. Six simulations were conducted at various initial NO (with δ15N=0 ‰) and O3 levels that were similar to
our experiment. Then the δ15N values of NO and NO2 during
the simulation were calculated from the model and are shown in Fig. B1,
suggesting that in our experimental condition, all systems should reach
isotopic equilibrium within 1 h.
Simulated NO–NO2 isotopic equilibrium process in the chamber at various NO and O3 concentrations.
Deriving Eqs. (7) and (8)
When the system (Reactions R1–R6) reaches a steady state, we have
d[15NO2]/dt=0.
Therefore, using Reactions (R1)–(R6)
k1[15NO2][14NO]+j(NO2)α1[15NO2]=k5α2[15NO][O3]+k1α(NO2–NO)[15NO][14NO2].
From here we refer to 14NO2 and 14NO as NO2 and NO for
convenience. Rearranging the above equation, we get
[15NO2][15NO]=k5α2O3+k1α(NO2–NO)[NO2]jNO2α1+k1[NO].
Meanwhile, since the Leighton cycle reaction still holds for the majority of isotopes (NO and NO2), we have
jNO2[NO2]=k5[NO][O3].
Thus,
[NO2][NO]=k5×O3jNO2.
From the text, when jNO2>0, we defined A=τexchange/τphoto=jNO2/(k1×[NO]). Using
the above equations, we know
C6jNO2[NO]=k5O3[NO2]=Ak1C7jNO2k1[NO]=k5O3k1[NO2]=A.
Next, to calculate δ(NO2)–δ(NO), we use the definition
of delta notation:
C8δ(NO2)–δ(NO)=RNO2/RSD-RNO/RSD=(RNO2/RNO-1)(1+δ(NO))C9RNO2RNO=15NO2[NO]15NO[NO2]=k5α2O3NO+k1α(NO2–NO)NO2[NO]jNO2α1[NO2]+k1NO[NO2].
Divide both sides by k1[NO][NO2]:
RNO2RNO=k5α2O3k1[NO2]+α(NO2–NO)jNO2α1k1[NO]+1.
Rearrange and substitute k5O3k1[NO2]
and jNO2k1[NO] with A:
C11RNO2RNO=α2A+α(NO2–NO)α1A+1C12RNORNO2=α1A+1α2A+α(NO2–NO)C13RNORNO2-1=α1-α2A-(α(NO2–NO)-1)α1A+α(NO2–NO).
Thus,
δ(NO2)-δ(NO)=α2-α1A+(α(NO2–NO)-1)α1A+α(NO2–NO)(1+δ(NO2)).
Then, using mass balance
δ(NO2)f(NO2)+δ(NO)(1-f(NO2))=δ(NOx)
we can derive Eq. (8):
δ(NO2)-δ(NOx)=α2-α1×A+α(NO2–NO)-1α1A+α(NO2–NO)1+δ(NO2)1-f(NO2).
Data availability
Data acquired from this study were deposited at Open Sciences Framework (Li, 2019; 10.17605/OSF.IO/JW8HU).
Author contributions
JL and GM designed the experiments, XZ and JL
conducted the experiments. XZ, GM, JO and GT
helped JL in interpreting the results. The manuscript was written by JL, and all the authors have contributed during the revision of the manuscript.
Competing interests
The authors declare that they have no conflict of interest.
Acknowledgements
We thank the NCAR's Advanced Study Program for providing support to Jianghanyang Li. The National Center for Atmospheric Research is operated by the University Corporation for Atmospheric Research under the sponsorship of the National Science Foundation.
Financial support
This research has been supported by the National Science Foundation.
Review statement
This paper was edited by Jan Kaiser and reviewed by Matthew Johnson and one anonymous referee.
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