Introduction
Hydrogen oxides (HOx= HO2+ OH) and nitrogen oxides
(NOx= NO2+ NO) play central roles in atmospheric
photochemistry. HO2, a product of OH-initiated volatile organic compound (VOC) oxidation, reacts
with NO to produce NO2, a key step in the photochemical ozone formation
cycle in the troposphere. Peroxynitric acid (often referred to as PNA,
HO2NO2, or HNO4) plays an important role in the coupling of
atmospheric HOx and NOx cycles (Niki et al.,
1977), especially at low temperatures. PNA serves as an important HOx
and NOx reservoir species altering the oxidative capacity of the
atmosphere on regional and global scales (Kim et al., 2007; Chen et al.,
2001; Davis et al., 2001; Carpenter et al., 2000).
PNA is formed via the reaction of HO2 and NO2 (DeMore et al.,
1997; Sander et al., 2011).
HO2+NO2⟶MHO2NO2
Formation via Reaction (R1) is favored at low temperatures and high pressures
(Kim et al., 2007). Unimolecular decomposition is temperature dependent and occurs with a lifetime of approximately 10 s, 1 atm and 298 K with the lifetime rapidly increasing to hours at 258 K
(Gierczak et al., 2005):
HO2NO2⟶MHO2+NO2.
PNA can be lost via photolysis in the near ultraviolet
(Jimenez et al., 2005) and near infrared (IR) via an
overtone band (Roehl et al., 2002; Stark et al., 2008):
HO2NO2⟶λOH+NO2,HO2NO2⟶λHO2+NO2,
or by reaction with OH (Jimenez et al., 2004):
HO2NO2+OH⟶H2O+NO2+O2.
In the lower troposphere, Reactions (R3), (R4), and (R5) typically occur on
timescales of days to months thereby implying that the dominant loss of PNA
in the lower troposphere is typically unimolecular dissociation or
deposition.
Deposition of PNA on snow surfaces has been observed in various studies
where the atmospheric lifetime of PNA in polar regions is largely controlled
by dry deposition (Huey et al., 2004; Slusher et al., 2002; Jones et al.,
2014). Additional laboratory studies have been performed confirming the
efficient uptake of PNA to ice (Li et al., 1996; Ulrich et al., 2012) and
sulfuric acid solutions (Zhang et al., 1997). Dependent on the
fate of PNA after deposition, the formation and subsequent deposition of PNA
has been suggested to result in a net loss of HOx and an increase in NO
(Grannas et al., 2007), with Reactions (R1)–(R5) thereby having an impact
on tropospheric ozone formation (Salawitch et al.,
2002).
Observations of PNA are generally limited in scope with most measurements
focusing on polar regions (Slusher et al., 2002, 2010; Huey et al., 2004), the free troposphere (Murphy et al., 2004; Singh et al.,
2006, 2007; Keim et al., 2008; Kim et al., 2007), and the
stratosphere (Rinsland et al., 1996, 1986; Sen et al.,
1998). Mean PNA observations from these studies range from tens of pptv in
polar surface regions to several hundred pptv in the upper troposphere–lower
stratosphere. The impacts of PNA on upper tropospheric chemistry have been
widely discussed (Brune et al., 1999; Wennberg et al., 1999; Faloona et
al., 2000), with one study in particular identifying a PNA contribution as
high as 20 % of the total NOy budget
(Murphy et al., 2004). Lower tropospheric,
mid-latitude measurements, in contrast, remain largely unexplored with the
exception of, to our knowledge, a single airborne study conducted in Mexico
(Spencer et al., 2009), where PNA concentrations up to 600 pptv were
observed and correlated with O3 formation.
The general lack of lower tropospheric, mid-latitude observations is driven
in part by two factors: (i) a diminished atmospheric impact of PNA due to
higher rates of thermal decomposition with respect to the generally colder
polar and upper atmosphere and (ii) a lack of instrumentation capable of
providing sensitive unambiguous measurements of PNA. Many of the techniques
available measure PNA as a component of NOy using O3/ NO
chemiluminescence (Keim et al., 2008) or total peroxynitrates via thermal
decomposition laser induced fluorescence
(Murphy et al., 2004). Currently, available
instrumentation capable of unambiguous measurement of PNA is limited to
remote sensing detection via IR absorption spectroscopy (Rinsland et al.,
1986, 1996; Sen et al., 1998) or in situ measurement via chemical
ionization mass spectrometry using the SiF6- ion (Slusher et
al., 2002, 2001; Huey, 2007), CF3O- (Spencer et
al., 2009; Huey et al., 1996), and iodide ion (I-) (Abida et
al., 2011). Among these, CIMS techniques have been shown to have sufficient
sensitivity and time resolution for the in situ monitoring of PNA
concentrations in the lower troposphere.
In this work, we present laboratory and ambient measurements illustrating
the utility of iodide ion CIMS for unambiguous measurement of PNA.
Additionally, we show applicability of this technique for the detection of
both HO2 and HONO, atmospheric species that are also integral to
HOx and NOx cycles. These results were necessary in order to rule
out potential mass overlap or PNA interferences from the sampling of
HO2 and HONO. A photo-source has been developed for dynamic production
of PNA, HO2, and HONO to assist with laboratory calibration and
elucidation of various ionization schemes. PNA observations made during the
2013 and 2014 Uintah Basin Wintertime Ozone Study (UBWOS) will be presented
and compared to the results of a chemically explicit box model developed to
describe the air quality in the Uintah Basin during a high ozone event
observed during the 2013 study (Edwards et al., 2014). The impact of PNA
on HOx and NOx budgets, particularly as it relates to the photochemical
production of ozone, will also be discussed.
Laboratory results
Laboratory experiments using I- CIMS were preformed both in preparation
and upon the conclusion of the UBWOS 2014 study. The goal of these
experiments was to adapt the I- CIMS technique for the sensitive
detection of PNA and develop a method for calibration. The development of an
HO2 based photolysis source for the production of HO2NO2 also
led to the recognition that the I- CIMS can be applied to the direct
measurement of HO2 radicals. Additionally, the HO2 photolysis
source is readily adaptable for the production of an online HONO
calibration standard. These laboratory developments in calibration standard
production and I- CIMS detection of PNA, HONO, and HO2 will be
discussed in detail in the following sections.
Standard generation and detection
The following describes the ion molecule reactions with I- ions
resulting in the detection of HO2NO2, HONO, and HO2:
I-+HO2⟶I⚫HO2-,I-+HO2NO2⟶I⚫HO2NO2-,I-+HO2NO2⟶HOI+NO3-,HO2NO2⟶ΔHO2+NO2⟶I-I⚫HO2-,I-+HONO⟶I⚫HONO-,
where Reaction (R7) is only observed using a cold inlet while Reaction (R9)
occurs upon thermal dissociation of PNA in the inlet. Experimentally, the
above reactions, with the exception of Reaction (R8), were observed in this
work to occur predominantly via reaction of the hydrated iodide cluster
(I–H2O-; m/z 145) based on the strong dependence of sensitivity on
water vapor observed during these experiments. This implies that the above
reactions are ligand switching reactions made faster by the ability of
H2O to accommodate excess energy of reaction through extra degrees of
freedom up to and including dissociation.
Shown are normalized (1e6 cps I-), background-corrected mass spectra acquired in the laboratory for calibration sources of
HO2 (a), HO2NO2 (b) and (c), and HONO (d). Spectra colored in
blue were collected using a room temperature inlet (∼ 25 ∘C) while spectra in red indicate that an inlet
dissociator at a temperature of 130 ∘C was used. Background
mass spectra have been subtracted from the displayed mass spectra to
highlight the m/z ions that are produced via the I- CIMS ion chemistry.
All three species are detectable at a unique m/z, when an inlet dissociator
is not used, allowing for simultaneous detection of HO2,
HO2NO2, and HONO.
Normalized (106 cps I-), background-corrected mass spectra are
shown in Fig. 1 for each of the sources. The ratio of m/z 145 (I–H2O-) to m/z 127 (I-) is displayed as a percentage on each
panel for reference. These spectra represent the result of the subtraction
of a normalized background mass spectrum from a sample spectrum. The
background correction method that was applied varies for each species and is
dependent on the sample matrix, which will be described separately in the
following sections. In all cases, the mass ranges from 126 to 128 and 144 to
146, corresponding to the I- (m/z 127) and I–H2O- (m/z 145)
ions, were removed to simplify interpretation of the mass spectra. These
spectra allow for the identification of impurities in the photolysis sources
used as well as demonstrating the ions that were used for the unambiguous
detection of each analyte. Each of these sources will be discussed in detail
in the following sections.
HO2 radical
HO2 radicals are generated in the laboratory via photolysis of H2O in
the presence of O2 (Dusanter et al.,
2008). A mixture of approximately 100 sccm N2 and 0.5 sccm O2 was bubbled through water and diluted into
a 5 standard L min-1 flow of N2. The mixture is subsequently passed into a PFA
photolysis cell and irradiated with a 185 nm Pen-Ray® lamp. The N2 dilution flow is produced using boil off from a
high-pressure liquid nitrogen Dewar to limit the amount of NOx and VOC
in the system. Detection of HO2 via I- CIMS occurs through direct
observation of the parent ion cluster (I–HO2); therefore, it is not
necessary to add CO in order to titrate OH as no measurement interference is
expected. Addition of trace amounts of CO was found to increase the
concentration of HO2 produced, though trace amounts of NO and NO2
from the steel cylinder mixture resulted in an increase in PNA and HONO
backgrounds with increasing CO.
Figure 1a shows the difference mass spectrum of the HO2 radical mixture
less the instrument background. In this case, the instrument background was
the ion signal measured prior to turning on the 185 nm lamp. It is clear
from the mass spectrum that there is only a single dominant peak observed at
m/z 160 (I–HO2-). A method for the quantitative calibration of the
HO2 radical source and the I- CIMS instrument for the detection of
HO2 radicals will be discussed in Sect. 3.1.4. The sensitivity of
I- CIMS to the detection of HO2 was determined to be a function of
the mixing ratio of water in the flow tube as well as the extent of
clustering/declustering in the system; e.g., reduction of the
I–H2O- : I- ratio due to higher E / N in the declustering region,
results in a lower observed sensitivity. Additional work to characterize the
effect of humidity on the detection efficiency is necessary to refine the
potential of this method for ambient monitoring of HO2.
The difficulty of quantitative sampling through an inlet is a significant
limitation to the measurement of radicals in the atmosphere. Laboratory
experiments were performed to probe the effect of inlet length on sampling
of HO2 radicals produced in N2. The results are shown in Fig. 2a, as the count rate at a given residence time normalized to the count rate
at the shortest residence time. This reaction is likely first order in
HO2 therefore a log-linear fit is the most appropriate representation
of the data; however, the data have been fit using a linear curve for
simplicity. HO2 is lost at a rate of approximately 0.60 s-1 in a 6 mm o.d. PFA inlet. Results indicate that the loss of HO2 is not driven
by surface losses, but by loss of HO2 via reaction with residual NO2
in the system to produce HO2NO2, as can be observed in the nearly
equivalent rate of increase in observed HO2NO2 (0.67 s-1,
Fig. 2b). Qualitative observations of H2O2, at m/z 161
(I–H2O2-), during the same experiment suggest that there is
no loss of HO2 via self-reaction occurring on these timescales.
Observed losses versus inlet residence times for a generated
standard of HO2 radicals sampled through PFA tubing at various flow
rates (3–6 standard L min-1) and lengths (0–3 m). Concentrations have been normalized to
the initial concentration observed at the minimum reaction time displayed.
Reaction with NO2 to form HO2NO2 appears to be the dominant
loss for HO2 on these timescales.
Peroxynitric acid
Two methods were used in this work for the production of a PNA standard. In
the first of these methods, PNA was synthesized using the techniques
described in Appelman and Gosztola (1995). Briefly, a nitrite-peroxide solution (NaNO2 in 30 %
H2O2) is mixed with a peroxide-acid solution (30 %
H2O2 in 70 % HClO4) at -20 ∘C to produce
approximately 1.7 M PNA in H2O2. The resulting solution is placed
in a glass diffusion cell (Williams et al.,
2000), at a temperature of -20 ∘C with zero air passed
over the headspace to produce a dynamic mixture of PNA. The 20 sccm
diffusion source outflow was sampled directly into the inlet flow of the
I- CIMS. This method of synthesis also results in the production of
non-negligible amounts of HNO3 and H2O2. While nylon wool can
be used to semi-selectively remove HNO3 from the calibration flow, no
method for the selective removal of H2O2 was identified. In any
case, HNO3 and H2O2 are observed at unique m/z ions and
therefore do not interfere with PNA measurement.
Alternatively, PNA can be dynamically generated using the output of the
HO2 source described in Sect. 2.2.1 (Ulrich
et al., 2012). Addition of NO2 to the output of the HO2 radical
source results in the production of PNA. Due to the relative simplicity of
this technique, photo-production of PNA was used as the preferred I-
CIMS calibration method for the laboratory and field measurements.
Figure 1 shows a difference mass spectrum of HO2NO2 detected using
an iodide CIMS instrument with a cold inlet (Fig. 1b) and an inlet
dissociator temperature of 130 ∘C (Fig. 1c). In both
cases, the instrument background was chosen as the ion signal prior to the
addition of NO2. When using a cold inlet, 30 ∘C, the dominant peak observed is m/z 62
(NO3-, Reaction R8). An ion signal at m/z 206
(I–HO2NO2-, Reaction R7) is also observed, although to a much lesser
extent than m/z 62. When an inlet dissociator
is used, HO2NO2 is observed at m/z 62 (NO3-) and m/z 160
(I–HO2-), where the detection of PNA at m/z 62 (NO3-)
results from incomplete thermal dissociation of HO2NO2 in the
inlet.
HONO
Similarly to PNA, HONO can be formed by addition of NO to the output of the
HO2 source described in Sect. 3.1.1. Addition of excess NO to the
HO2 calibration source results in the production of HONO from titration
of HO2, as well as any OH produced in the source, via the following
reactions:
HO2+NO⟶OH+NO2,OH+NO⟶HONO.
Figure 1d shows the difference mass spectrum of the HONO calibration source,
where the instrument background here was chosen as the ion signal prior to
the addition of NO. It is clear from the figure that HONO is the only
product formed and is detected by I- CIMS at m/z 174 (I–HONO-).
This result is in contrast to the Abida et al. (2011) study which also reports
m/z 46 as a minor ion, a difference that can be attributed to the relatively
stronger clustering used in our work.
This method of HONO production is instantaneous and does not require the
period of stabilization that is necessary for acid–salt reaction-based
sources (Febo et al., 1995). HONO standard production via
the reaction of HO2 and NO provides a good alternative to previously
used I- CIMS calibration methods (Roberts et al., 2010).
Dynamic source calibration
Quantification of PNA and HONO produced using the above-described methods
was performed using the quartz catalysis total NOy instrument
(Wild et al., 2014) described in Sect. 2.2.2.
Laboratory experiments indicate that more than 99 % of PNA is thermally
dissociated above a temperature of 100 ∘C while HONO
decomposition is negligible below 200 ∘C. The quartz inlet
was operated at gas temperatures of 160 and
720 ∘C for the measurement of PNA and HONO, respectively.
The difference in total NOy minus the sum of NO2 and NO detected
yields a quantitative measurement of the PNA or HONO produced in the source.
During these experiments the calibration source flow was alternately sampled
by the I- CIMS and CRD instruments. In order to eliminate any
differences in radical reaction times, as a result of inconsistencies in the
inlet lengths between the two instruments, the gaseous mixture was passed
over glass wool subsequent to addition of NO or NO2 to terminate the
reaction by removing any remaining HO2 radicals. In this manner the
I- CIMS sensitivity is calculated as the ratio of the I- CIMS ion
signal to the cavity ring-down spectroscopy (CaRDS) measured concentrations. A summary of instrument
sensitivities and detections limits (3σ) is included in Table 1. For
the calibration data reported in Table 1, the m/z 145 to m/z 127 ratio was
approximately 30 %.
Summary of observed products, sensitivities, and detection limits
(DL) for the reaction of I- with HO2NO2, HONO, and
HO2.
Analyte
Detected Ion
Sensitivitya
DL
Inleta
(m/z)
(Hz pptv-1)
(pptv, 3σ)
HO2NO2
I–HO2NO2- (206)
0.40 ± 0.06
20
Cold
NO3- (62)
144 ±11b
0.7
Cold, Hot
I–HO2- (160)
2.0 ± 0.04
40
Hot
HONO
I–HONO- (174)
1.7 ± 0.3
30
Cold, Hot
HO2
I–HO2- (160)
2.6 ± 0.3
20
Cold, Hot
a Indicates the inlet dissociation temperature at which the particular
ion can be sensitively observed. The inlet dissociator was operated at a
temperature of 130 ∘C (hot) or at ambient temperature
(cold).
b Sensitivity reported is for detection with an inlet dissociator at
ambient temperature (cold).
The I- CIMS sensitivity towards HONO at the I–HONO- cluster (m/z 174), Reaction (R10), was determined to be 1.7 Hz ppbv-1 with a
corresponding 3σ instrumental detection limit of 30 pptv. The
detection limit for HONO is largely limited by the magnitude of the
ever-present I–NO2- (m/z 173) ion and the low resolution of the
quadrupole mass spectrometer used in this work. The HONO sensitivity
reported here represents detection with a 25 ∘C inlet;
however, the instrument sensitivity was found to be nearly equivalent,
within the stated uncertainties, using the heated (130 ∘C)
dissociator.
The sensitivity towards PNA was determined for detection at the
I–HO2NO2- (m/z 206, 25 ∘C inlet),
I–HO2- (m/z 160, 130 ∘C inlet), and NO3-
(m/z 62, 25 ∘C inlet) ions, Reactions (R7), (R6), and (R8),
respectively. The most sensitive method of detection was observed via NO3-
in a 25 ∘C inlet, 144 Hz pptv-1with a corresponding
3σ detection limit of 0.7 pptv. Detection at the I–HO2-
and I–HO2NO2- ions are considerably less sensitive, 2.0 and
0.4 Hz pptv-1, respectively. The 3σ detection limits are
approximately 40 and 20 pptv for I–HO2- and
I–HO2NO2-, respectively.
While the most sensitive detection of PNA occurs via NO3-, there exist
several potential interferences that are also observed at that ion; see
Table 1 in Wang et al. (2014). Considering the
sensitivity to PNA detection, relatively low daytime levels of PNA will
result in significant signals at m/z 62. In fact, a recent study has
suggested that evidence exists for a large daytime source of N2O5
detected via NO3- ion using I- CIMS measurements (Wang et al.,
2014); however, low daytime levels of HO2NO2 would also be
consistent with observations presented in that work. Unfortunately,
detection of PNA at m/z 160 (I–HO2-) leads to overlap with
HO2 radical detection at that mass. Therefore, we suggest that m/z 206
(I–HO2NO2-) will yield the most reliable, interference free
method of PNA detection for ambient measurements. Increasing the I–H2O
cluster ratio beyond 30 % (this study) should improve the instrument
sensitivity towards detection of PNA at the m/z 206 cluster ion.
An indirect calibration was performed for the quantification of the HO2
source, as no direct HO2 measurement was readily available. An initial
amount of HO2 radicals are generated and monitored via the I- CIMS
operated with a 25 ∘C inlet. A small amount of NO2 is
then added generating PNA with a corresponding reduction in observed I-
CIMS HO2 signal. The concentration of HO2 lost by titration is
assumed to be equivalent to the amount of PNA produced in the reaction, as
the formation of PNA involves the consumption of one HO2 radical per
molecule. Using the previously described PNA calibration method, the
sensitivity can be calculated as the ratio of the reduction in the observed
I- CIMS I–HO2- signal to the NOy measured PNA
concentration. The I- CIMS detection sensitivity was determined using
this method to be 2.6 Hz pptv-1 with a corresponding 3σ
detection limit of 20 pptv. As it was not the focus of this work, the
instrument inlet was not optimized for the sampling of radical species;
therefore, changes in the inlet design and optimization of the iodide–water
cluster distribution in the flow tube could both serve to increase the
instrument sensitivity and improve instrument detection limits to levels
more appropriate for ambient sampling of HO2.
Uintah Basin Wintertime Ozone Study observations
A comparison of HO2NO2 observations made using the
I–HO2- (m/z 160) ion with a hot dissociator (130 ∘C) and the I–HO2NO2- (m/z 206) ion with a cold dissociator
(∼ 25 ∘C). Inset is a correlation plot of the
two measurements where comparison is possible.
Observations of HO2NO2 during the UBWOS 2014 study are shown in
Fig. 3 for the entire duration of the measurement period. Measurements of
the I–HO2NO2- ion using a cold inlet were performed only
during the initial and final portion of the study for 30 minutes every
hour. Ambient air was sampled through a heated inlet dissociator on
alternating 30 min periods with cold sampling performed on the opposite
time periods during the following: 24–30 January and 4–14 February. For the
remainder of the measurements, 30 January–4 February, ambient air was
continuously sampled through a heated dissociator. During sampling periods
where the inlet dissociator was used, PNA was monitored on m/z 160
(I–HO2), and during periods of cold sampling via m/z 206
(I–HO2NO2). It is clear from the correlation plot inset in Fig. 3 that both sampling methods (cold and hot) agree reasonably well; slope
= 0.93; R2= 0.785. There does appear to be a positive bias in the
m/z 160 observations relative to detection at m/z 206, indicated by a
positive intercept and clear disagreement during certain periods shown in
the Fig. 3 time series. One possible explanation is the sensitivity at m/z 160 to ambient HO2 or HO2 generated in the inlet as a product of
PNA decomposition. Given the length, temperature, and residence time of the
inlet used (20 m, 30 ∘C, ∼ 4.8 s), PNA is
expected to decompose by approximately 5 % prior to sampling. The HO2
radicals produced as a result of PNA decomposition would likely be detected
via m/z 160 yet remain unobserved on the m/z 206 ion leading to a positive bias
in m/z 160 observations. Additionally, from laboratory results presented in
Fig. 2b, we expect there to be production of HO2NO2 in the
inlet; however, model approximations of the HO2 mixing ratio during the
2013 study (15 pptv maximum) yields a maximum formation of HO2NO2
in the inlet during peak ozone events of ∼ 1 % (Edwards
et al., 2014).
Aside from the sampling conditions already described, several days of PNA
measurements were performed during the 2014 study comparing a short unheated
PFA inlet (3 m, ambient temperature) and a longer heated inlet (20 m,
30 ∘C). An average reduction of 5 % is observed in the PNA
mixing ratio when sampling through the longer heated inlet as a result of
thermal and perhaps surface assisted decomposition. It is important to note
that the data presented here are not corrected for these inlet losses and
are therefore considered a lower limit on ambient PNA.
During UBWOS 2013, the I- CIMS was equipped to monitor PAN compounds
using a thermal dissociation inlet (150 ∘C). This was done
prior to the laboratory work described in this manuscript and neither m/z 160 nor m/z 206 were monitored during the measurement period. However, the
NO3- ion (m/z 62) was monitored throughout the entire campaign
with observed signals exceeding 2e5 counts per second at times. Using the
calibration data obtained from the laboratory portion of this work, an
approximate PNA concentration was determined for the 2013 UBWOS study,
assuming the same NO3- (m/z 62) sensitivity as measured during the
laboratory calibrations performed after the 2014 study (144 Hz pptv-1),
corrected for differences in dissociator temperatures and transmission of
HO2NO2 through the heated inlet tip used during the 2013 study.
The error associated with this method was calculated to be approximately
60 %, largely due to corrections applied to account for differences in
instrument tuning and additional HO2NO2 losses due to the
different inlet conditions used during the 2 years.
Diurnal profiles of PNA are shown in panel (a) for the full
measurement period during the UBWOS 2013 and 2014 studies, with the 2014
study separated based on the sampling height location. Panel (b) presents
the difference in the 1 and 18.4 m PNA measurements for duration of the 2014
study, where a positive value (brown) indicates larger concentrations at the
ground and a negative value (blue) suggests a relative PNA depletion in the
1 m measurements. The shaded region represents 1 standard deviation on the
hourly average for the entire measurement period.
Figure 4a shows the diurnal average of 2013 I- CIMS observations of PNA
for the entire study. While the average diurnal mixing ratio peaks at 0.5 ppbv, mixing ratios up to 1.5 ppbv were observed during the 2013 study and
can be explained by the coincidence of high daytime levels of NO2 with
the low temperatures in the Uintah Basin. Similarly to the 2014
measurements, also included in Fig. 4a, PNA reaches a peak after solar
maximum (∼ 15:00 MST) with a minimum observed throughout the
night. Concentrations of HO2NO2 observed during the 2014 study
were significantly lower relative to the 2013 study with a maximum average
mixing ratio of 0.1 ppbv. During the 2013 and 2014 study, N2O5, a
nighttime species and potential interference on the NO3- ion, was
not observed to contribute to the observed daytime signal.
Temperature and mixing ratios of HO2 and NO2 required to
sustain an equilibrium concentration of 1 ppbv HO2NO2. The region
within the dashed circle superimposed on the figure highlights the
conditions encountered during the 2013 UBWOS study. Data shown were
calculated using the IUPAC database (Atkinson et al.,
2004).
The PNA mixing ratios observed during the 2013 and 2014 studies are expected
for cold conditions with active photochemistry and sufficient NOx
pollution. Displayed in Fig. 5 are the conditions necessary to sustain an
equilibrium concentration of 1 ppbv PNA with respect to temperature and the
mixing ratios of HO2 and NO2. The dashed area superimposed on the
figure represents the approximate range of conditions encountered during the
2013 study, where HO2 levels were approximated using model results that
describe an ozone event observed during the 2013 study, described below. The
Uintah basin provided a unique atmosphere that promotes the formation of PNA
for several reasons (1) a strong inversion during the wintertime allows for
concentrations to build up in the boundary layer over several day periods
(2) low ambient temperatures favoring the formation of HO2NO2 over thermal decomposition and (3) radical species propagation,
e.g., HO2 formation, is enhanced due to the active chemistry observed
during ozone formation events (Edwards et al., 2014). Lastly, as will be
discussed later, the snow surface acts as an important interface serving as
both a source and a sink of HO2NO2.
The conditions encountered in the basin between the 2013 and 2014 season can
be used to explain the large difference in the observed PNA ambient mixing
ratios, shown in Fig. 4a. Mainly during the 2013 study we observed strong
inversions over multi-day periods allowing for the buildup of primary and
secondary pollutants, a phenomenon driven by snow surface cover and
meteorological conditions encountered during the 2013 study (Ahmadov et
al., 2015; Edwards et al., 2014). In contrast, relatively low snow cover
during the 2014 season limited the formation of multi-day inversions
yielding lower ambient mixing ratios of both primary and secondary
pollutants. The resulting combination of lower NOx mixing ratios and
higher ambient temperatures during the 2014 study thereby favored thermal
dissociation of HO2NO2 and led to lower ambient mixing ratios than
observed during the 2013 study. In addition, limited snow cover and reduced
deposition of NOy species to the snow surface, as a result of lower
ambient mixing ratios, likely reduced the role of the snow surface as a
source of HO2NO2 in 2014.
As described previously, the vertical gradient of various species was probed
through the use of a dual inlet system (18.4 and 1 m heights) during the
UBWOS 2014 study. A comparison of those measurements for PNA is shown in
Fig. 4a as diurnal averages of the measurements where dual inlet switching
was applied. Figure 4b shows the result of the difference of the 1m minus
the 18.4 m PNA measurements where the shaded region represents 1 standard
deviation of the average for the entire study. Displayed in this fashion, a
positive value is an indication that PNA is larger in the surface coupled
layer, characteristic of daytime observations, than aloft with the opposite
indicating a relative enhancement of PNA in the layer decoupled from the
surface, which was typically observed at night. On average, HO2NO2
is depleted at the surface relative to air aloft by approximately 15 pptv
with a reversal of nearly the same magnitude observed during mid-day. It is
important to note that this data represent a 1 h average over the entire
6-week measurement period, and while the overall magnitude shown here is
small, observed ΔPNA values ranged from -150 to 150 pptv.
One interpretation of these results is that deposition of PNA to the snow
surface occurs throughout the night in the Uintah Basin with an emission
from the snow surface observed in the early morning to mid-day. Through this
mechanism snow surface photolysis of PNA could serve as an additional
daytime surface source of NO2 and HO2. A similar result has been
reported by a recent study (Jones et al., 2014),
where evidence for surface exchange was observed for HO2NO2 and
HNO3 during the Antarctic winter. However, previous studied have
highlighted the complexities of photochemistry chemistry occurring at or
directly above snow surfaces (Chen et al., 2007, 2001), which
complicates the interpretation of these observations.
Specifically, previous studies have shown enhancements of reactive species
integral to HOx and NOx budgets at snow surfaces, e.g., HONO,
CH2O, and H2O2 (Hutterli et al., 1999,
2001; Honrath et al., 1999; Ridley et al., 2000; Zhou et al., 2001). Photolysis
of these HOx/ NOx precursors at or above the snow surface or a
direct emission of NOx could result in a net apparent surface source of
HO2NO2, as it would shift the gas phase equilibrium, Reactions (R1) and (R2),
towards the formation of HO2NO2. The net
effect of this process on the HOx budget at the snow surface is
difficult to quantify without simultaneous observations of the other
dominant HOx precursors and an understanding of their chemical fates.
Therefore, further validation of our interpretation, a bidirectional flux of
HO2NO2 from the snow surface, awaits direct flux measurements of
HO2NO2 over these highly polluted snow surfaces.
Regardless of the underlying mechanisms, the dynamics causing the observed
vertical distribution in PNA observations, whether mixing or deposition in
origin, can have a measurable impact on the ozone formation potential in
the Uintah Basin. To investigate the effects, a chemical box model, based on the Master Chemical Mechanism (MCM) v3.2 chemistry scheme, has been developed to describe observed ozone
production during a wintertime ozone pollution episode during UBWOS 2013
(Edwards et al., 2014). This 0-D model contains a near-explicit oxidation
mechanism for 32 observed VOC and oxidized VOC, and is constrained using
constant emissions of primary species (VOC and NO), with the concentrations
of all other species calculated by the chemistry scheme. Physical loss
processes, such as mixing and deposition, are represented via a bimodal
first-order loss process for all species, the rate of which has been
optimized based on boundary layer height observations and the concentrations
of long-lived species, e.g., methane.
Figure 6a and b display the base model calculation of PNA and ozone
reported in the Edwards et al. (2014) study. While there is reasonable
quantitative agreement between the model and measurement daily maxima, there
is a temporal shift in PNA measurements relative to model predictions. While
the reasons for this are as of yet unknown, possible explanations include
underestimation in modeled NO, additional daytime PNA loss mechanisms, or
issues with the simple parameterization of mixing used in the model. The
latter seems unlikely as the model reproduces the observed diurnal variation
in ozone relatively well. As discussed above, there is also evidence that
PNA is lost to the snow surface, though the temporal trend in deposition
implied from the gradient measurements does not suggest a relatively higher
rate in the morning to early afternoon than in the evening, rather there is
a possible indication of a mid-day snow surface source. Nitrogen oxides
within the model are parameterized using a constant source of NO, with the
partitioning of all nitrogen oxides calculated by the chemistry scheme. A
quantitative comparison of NO observations to modeled values suggests that
the model does typically underestimate the NO concentration throughout the
morning. This underestimation would lead to an underestimation of the loss
of HO2 via reaction with NO, thereby slowing the formation rate leading to
an overprediction of PNA. This process would have a particular impact in
the morning hours where observed NO is relatively large and the HO2
source is limited.
Comparison of PNA and ozone observations throughout an ozone
formation event observed during UBWOS 2013 and corresponding model
predictions using an explicit chemical box model describing the chemistry.
Panel (a), HO2NO2 observations compared to model results
applying various PNA lifetimes with respect to deposition (tPNA). The
datum shown with error bars, black circle, represents the approximated
60 % error on the 2013 HO2NO2 I- CIMS measurements. Panel
(b) illustrates the effect varying deposition rates of PNA has on total
predicted ozone production.
Deposition of PNA has the potential to result in a net loss of HOx and
thus an increase in NO (Grannas et al., 2007), due to reduced titration
to NO2, which would have an overall effect on ozone formation
potential. The sensitivity of ozone production in the UBWOS 2013 box model
to changes in the lifetime of PNA with respect to deposition (tPNA) was
investigated and the results of these tests are displayed in Fig. 6. In
addition to the first-order physical loss term applied to all species,
predominantly representing losses due to mixing (Edwards et al., 2014),
an additional first-order loss term for PNA was added to the model scheme
to represent deposition. Calculations with a lifetime for PNA with respect
to deposition (tPNA) of 1 h yield approximately a 12 % reduction
in daily max ozone from the base case where no PNA deposition was included.
It is not possible to determine the observed lifetime of PNA with respect to
deposition to the snow surface using data from the UBWOS 2013 or 2014
studies; however, the measurements and model suggest that surface deposition
is occurring with a potentially measurable effect on the ozone production
in the Uintah Basin.
The post-depositional fate of PNA is also important as this scenario assumes
that the loss of PNA to the snow surface is irreversible and thus a net
HOx and NOx sink. Subsequent snow chemistry resulting in storage
and volatilization of HOx (Jones et al.,
2014) or NOx, has the potential to reduce the magnitude of the
effects observed in Fig. 6b. If PNA dissolves in an aqueous solution, such
as a quasi-liquid layer on the snow surface in the Uintah basin, it can
undergo the following dissociation and ionic reactions (Logager and
Sehested, 1993; Zhu et al., 1993; Goldstein et al., 2005):
HO2NO2⟷HO2NO2(s),HO2NO2⟷pKa∼6H++O2NO2-,HO2NO2(s)⟶HONO,O2NO2-⟶NO2-+O2,O2NO2-⟶NO2+O2-,HONO⟷pKa∼6H++NO2-.
Previous work has also shown that uptake of PNA to highly acidic surfaces
results in reversible uptake, where chemical loss to NO2- is
negligible (Zhang et al., 1997).
Displayed are gas phase HO2NO2 observations made during
two wintertime ozone events in the Uintah Basin, UT, observed in 2013 (a),
and through the entire 2014 study (b). Measurements of the nitrite content
of the snow surface taken during the same time periods are also shown. In
2013, a precipitation event that occurred on the afternoon of 8 February, shown as the period in blue, added fresh surface snow and flushed pollutants
out of the basin resulting in lower ambient PNA and snow nitrite levels.
Panel (c) shows in more detail periods from the 2013 measurements to better
illustrate the daily reduction in snow surface nitrate that was regularly
observed.
Measurements of snow surface nitrite made during the 2013 and 2014 study are
shown in Fig. 7a and b, respectively, along with observations of PNA.
Two pollution events observed in 2013 were separated by a cleanout event
where additional precipitation (indicated by the blue shaded region in
Fig. 7a) accompanied by higher winds and unstable conditions ventilated
the basin. The nitrite content of the snow surface generally increases
throughout the first event at a rate proportional to the daily increase in
PNA mixing ratio until fresh snow is added and nitrite levels drop.
Throughout the second event, nitrite levels in the snow surface again build
up as ambient PNA levels also increase. We note here that any
HO2NO2 dissolved in the snow can form NO2- with
efficiencies as high as 56 % depending on pH (Goldstein
et al., 2005).
A more detailed look at snow nitrite content shows a generally observed
decrease in surface layer concentrations throughout the day, Fig. 7c, a
result that is consistent with reversible uptake of nitrite. This daytime
depletion of snow nitrite could be evidence of the reversible uptake of PNA,
interpreted from data shown in Fig. 4a, or an indication of the formation
and subsequent release of other nitrogen-containing species, such as HONO or
NOx, from the snow surface. It is interesting to note that the highest
measurements of snow nitrite in a given day typically occur prior to the
buildup of ambient PNA. This result is consistent with nocturnal uptake of
PNA or HONO to the snow surface, as shown in Fig. 4b; however, no evidence
of this nighttime enrichment is available, as nighttime measurements of snow
surface nitrite were not made. Additionally, the extent to which Reactions
(R13) and (R14) were occurring, and therefore PNA contributing to snow nitrite
content, cannot be approximated as the pH of the snow surface was not
measured during this study. Furthermore, while the deposition and potential
volatilization of HO2NO2 can contribute to the net flux of
nitrite at the snow surface, other species such as HNO3 and HONO are
also known to deposit to snow surfaces and should also be considered as part
of the net surface nitrogen budget.
Improved measurements of the deposition velocity on snow surfaces of varying
acidities and the chemical fate of PNA after deposition is necessary to
improve our understanding of the impacts of PNA deposition to HOx and
NOx budgets. As our model results indicate the post-depositional fate
of PNA can have a non-negligible impact on ozone formation potential,
particularly in cold regions, such as the Uintah Basin, where the lifetime
of PNA is sufficiently long such that deposition becomes a dominant sink.
Considering the similar surface chemistry and influence on radical budgets
for atmospheric HONO (VandenBoer et al., 2015),
simultaneous measurement must be performed to understand the tropospheric
fate of PNA and HONO species in cold regions and the extent of their
involvement in tropospheric HOx and NOx budgets.