Introduction
Schematic overview depicting processes that contribute to the
production of Oxygen emission bands in the middle atmosphere. The black solid
arrows are the most important gas phase reactions (k) and quenching
reactions (qi). The blue and green solid arrows are photodissociation and
photoexcitation processes respectively (Ji). Dashed arrows correspond
to spontaneous radiative emissions (Ai), and the red dashed arrows show
the radiative emissions that are the subject of this study. (*) in
O2* denotes several excited states (see text for explanation).
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The atmospheric airglow in the mesosphere and thermosphere above
≈ 60 km is formed by fluorescent emission from excited states of
atoms and molecules. Atoms and molecules in the mesosphere and lower
thermosphere (MLT) can be excited by absorption of solar radiation
(photoluminescence) or by exothermic chemical reactions (chemiluminescence; see, e.g., ).
The processes that contribute to the airglow of atomic and molecular oxygen
in the mesosphere and lower thermosphere are shown in Fig. .
The recombination of atomic oxygen , denoted as
Reaction () and (k) in Fig. , produces
O2* excited molecules:
O(3P)+O(3P)+M⟶O2*+M.
Here, O2* represents any of the seven states below the first dissociation
limit. Bates and others argue that the population distribution between these
states can best be approximated statistically, in which the
5Πg state is produced in almost 40 % of the collisions
. Most of the O2* derived
from recombination is found in the A3Σu+ state
, and in a recent review Huestis concludes that all of the
recombining atoms pass through the Herzberg states c1Σu-,
A′3Δu, and A3Σu+ .
provide the quenching parameters resulted from
analyzing the measurements of the near-ultraviolet portion of the nightglow
to fit the synthetic spectra of the Herzberg bands of O2. These
parameters set an upper limit of 10 % production efficiency on the
generation of O2(c1Σu-) in the atomic oxygen
association reaction . Admittedly, proper accounting of the correct products
of Reaction () can be complex. Recent research has
investigated this issue (e.g., ). Therefore
we assume the production of a surrogate “hybrid” state O2* in the
photochemical model.
The photolysis of ozone (JH in Fig. ,
Reaction ) in the Hartley band (λ<310 nm) leads to the
first electronically excited state of atomic oxygen O(1D) and
molecular oxygen O2(a1Δg) :
O3+hν(λ≤310nm)→O(1D)+O2(a1Δg).
The photolysis of molecular oxygen in the Schumann–Runge continuum
(JSRC in Fig. , Reaction ) and at
Lyman α (JLα in Fig. ,
Reaction ) leads to electronically excited oxygen atoms
O(1D) :
O2(X3Σg-)+hν(130≤λ≤175nm)→O(3P)+O(1D),O2(X3Σg-)+hν(λ=121.6nm)→O(3P)+O(1D).
Quenching (collisional de-excitation) processes are represented by black
solid arrows and denoted by (qi) in Fig. . The O2*,
produced by Reaction (), can be quenched by atomic oxygen
to produce O(1S) via Reaction () (q5 in
Fig. ) or quenched by molecular oxygen to
produce O2(b1Σg+) via
Reaction () (q6 in Fig. )
:
O2*+O(3P)→O(1S)+O2(X3Σg-),(q5)O2*+O2(X3Σg-)→O2(b1Σg+)+O2(X3Σg-).(q6)
O(1D), by quenching with O2(X3Σg-),
produces O2(b1Σg+) via Reaction (),
which combines q1 and q4 in Fig. :
O(1D)+O2(X3Σg-)→O2(b1Σg+)+O(3P).
Photoabsorption of the solar radiation at 761.9 nm produces
O2(b1Σg+) directly via
O2(X3Σg-)+hν(λ=761.9nm)→O2(b1Σg+)
in the sunlit mesosphere , shown in Fig.
as the radiative excitation J1.
Then, according to Reaction (), O2(b1Σg+)
can be reduced in energy to O2(a1Δg) by collisions
with any of the abundant species such as O2, N2, CO2,
or O (denoted by “M”), shown as q2 in Fig.
:
O2(b1Σg+)+M→O2(a1Δg)+M.
Note that the above is a fast process (spin-conserved).
O2(a1Δg) can in turn be quenched via
Reaction () (q3 in Fig. ) to produce
O2(X3Σg-):
O2(a1Δg)+M→O2(X3Σg-)+M,
which is a slow process (spin forbidden because the ground state is
O2(X3Σg-)).
Spontaneous radiative emissions are represented by dashed arrows and denoted
by (Ai) in Fig. . O(1S) decays to O(1D) by
emitting 557.7 nm photons (the oxygen green line, Reaction , A4
in Fig. ), which is fast (because it conserves the spin)
:
O(1S)→O(1D)+hν(λ=557.7nm).
Reactions (), (), and () are commonly
referred to as the Barth mechanism (; see, e.g., the
review by ). The green line emission allows the deduction of the
atomic oxygen densities near 100 km, as shown for example by
. The oxygen 297.2 nm line is one of the prominent
components of the ultraviolet nightglow , and it is
produced by O(1S) via Reaction () :
O(1S)→O(3P)+hν(λ=297.2nm),
which is indicated by A5 in Fig. .
O(1D) – produced from Reactions (, JH),
(, JSRC), (, Jα), or
(, A4) – can be deactivated to the ground state O(3P) by
the slow (spin-forbidden) emission:
O(1D)→O(3P)+hν(λ=630.0nm),
where the 630.0 nm red line is represented by A1 in
Fig. .
Among the strongest features of the day and night airglow are the infrared
atmospheric band (a1Δg→X3Σg-)
and the atmospheric band (b1Σg+→X3Σg-) of molecular oxygen . These two
spontaneous radiative emissions, which we deal with in this work, are
represented by the thick red dashed arrows in Fig. . They are
emitted by the deactivation of the two excited states of the molecular oxygen
O2(b1Σg+) at 761.9 nm via
Reaction () (A2 in Fig. ) and
O2(a1Δg) at 1.27 µm via
Reaction () (A3 in
Fig. ) :
O2(b1Σg+)→O2(X3Σg-)+hν(λ=761.9nm),O2(a1Δg)→O2(X3Σg-)+hν(λ=1.27µm).
Assuming that the processes in Fig. describe the
photochemistry and chemistry, one can deduce ozone densities from
measurements of the infrared atmospheric volume emission rates (VERs) in the
O2(b1Σg+) and O2(a1Δg)
bands (hereafter O2(1Σ) and O2(1Δ) bands
respectively). For this, the rates of all of these processes such as q1
and q4 in Fig. and in the Reactions ()
and () should be known
e.g.,.
Previous measurements
The oxygen airglow was measured from spaceborne platforms and rocket
experiments in several previous studies. Measurements of the
O2(1Σ) band include the Fabry–Pérot interferometer on the
Dynamics Explorer 2 (DE-2) satellite , which were used to
study the overall brightness of the emission. The high-resolution Doppler
imager (HRDI) on the Upper Atmosphere Research Satellite (UARS)
measured the Doppler shifts of rotational lines of the
O2(1Σ) atmospheric band to determine the winds in the
stratosphere, mesosphere, and lower thermosphere. The Wind Imaging
Interferometer (WINDII) on the same satellite measured
wind, temperature, and emission rates. The TIMED Doppler Interferometer
(TIDI) on the Thermosphere–Ionosphere–Mesosphere Energetics and Dynamics
(TIMED) satellite performed remote sensing measurements
of upper atmosphere winds and temperatures based on O2(1Σ)
emission. The Remote Atmospheric and Ionospheric Detection System (RAIDS) on
the International Space Station's Kibo module
measured the limb brightness of the O2(1Σ) (0,0), (0,1), and
(1,1) vibrational band emissions from 80 to 180 km. The Optical Spectrograph
and InfraRed Imaging System (OSIRIS), onboard the Odin satellite
, was used to derive temperatures in the mesosphere and lower
thermosphere region (MLT) from O2(1Σ).
Previous measurements of the O2(1Δ) band include observations
from the near-infrared spectrometer experiment on the Solar Mesosphere
Explorer satellite (SME). SME measured emission from O2(1Δ)
produced by photolysis of O3 . The infrared
atmospheric-band airglow radiometer (IRA) aboard the satellite OHZORA
measured the mesospheric ozone profile derived from O2(1Δ)
emission . One part of the Optical Spectrograph
and InfraRed Imager System instrument onboard the Odin satellite is a
three-channel infrared imager (IRI)
that observes the scattered sunlight and the airglow from the oxygen infrared
atmospheric band at 1.27 µm . The TIMED–SABER
(Sounding of the Atmosphere using Broadband Emission Radiometry) data were
used to measure the O2(1Δ) airglow emission by a channel with
central wavelength of 1.27 µm .
All of the abovementioned studies include satellite observations of only one
of the O2 bands, either O2(1Σ) or O2(1Δ).
Simultaneous measurements of both O2(1Δ) and
O2(1Σ) airglow were part of the Mesosphere–Thermosphere
Emissions for Ozone Remote Sensing (METEORS) sounding rocket experiment. It
was launched from White Sands Missile Range, New Mexico
, and was used to derive ozone concentrations
separately from each of the O2 bands.
Present work
In this work, we retrieve volume emission rates from the airglow of
the O2(1Σ) and O2(1Δ) bands in the
mesosphere and lower thermosphere from the SCanning Imaging Absorption SpectroMeter for Atmospheric CHartographY
(SCIAMACHY , and references therein)
onboard the European Space Agency Envisat satellite. We present the
retrieval algorithm and O2(1Σ) and O2(1Δ) band
volume emission rates. We analyze daily mean latitudinal distributions of
VERs in the altitude range of approximately 50–150 km.
In Sect. we describe the SCIAMACHY dataset and our
method to retrieve both the O2(1Δ) and the O2(1Σ)
volume emission rates. Results are presented in Sect. ,
including the retrieved volume emission rates and first results on the
temporal and spatial variations of the volume emission rates. We also include
one example study of the relation between the temporal variations and the
sudden stratospheric warming (SSW) in 2009. In Sect. we summarize
the findings of our study and give conclusions.
Data and methods
Data
The SCanning Imaging Absorption SpectroMeter for Atmospheric CHartographY
is a passive remote sensing spectrometer that observes back-scattered, reflected, transmitted, or emitted radiation from the atmosphere
and the Earth's surface in the 240–2380 nm wavelength range. The instrument
is part of the atmospheric chemistry payload onboard the Envisat satellite,
which was operational from March 2002 until April 2012. SCIAMACHY has three
different viewing geometries: nadir, limb, and moon–sun occultations. From
July 2008 until April 2012, SCIAMACHY observed the mesosphere and lower
thermosphere region (50–150 km) regularly twice a month. This special
MLT limb mode scans the mesosphere and lower thermosphere in 30 limb points
from 50 to 150 km altitude with a vertical spacing of about 3 km. These
scans were scheduled in place of the nominal mode scans and there were
20 limb scans along one semi-orbit. Overall 84 days of mesosphere and lower
thermosphere limb measurements were carried out. In this work, we use the
visible and near-infrared spectra from channel 4 (595–811 nm) and channel 6
(1200–1360 nm) in the MLT limb-viewing geometry to retrieve volume emission
rates (VERs) from the airglow of the O2(1Σ) and
O2(1Δ) bands. Use of the additional channels covering the green
line or UV is beyond our current work, and we refer to
.
To generate data for our study, we used the SCIAMACHY dataset level 1b
version 8.02 and the SCIAMACHY command line tool
SciaL1c
from the SCIAMACHY calibration tools. We selected two windows for each of the
two bands: 750–780 nm for the O2(1Σ) band (759–767 nm) and
1200–1360 nm for the O2(1Δ) band (1255–1285 nm). We subtract the
spectrum measured at ≈ 360 km tangent height as a dark spectrum
from the measured spectra at all of the other tangent heights. This spectrum
contains some residual spectral (readout) patterns left from the calibration
step, and subtracting it from other spectra which have almost the same
patterns cancels out that. For the O2(1Δ) band, there are two
masked points that appear in every scan located around 1262 and 1282 nm.
Daytime spectra
A typical orbit starts with a limb measurement of the twilit atmosphere,
followed by the solar occultation measurement during sunrise over the North
Pole and an optimized limb-nadir sequence . Our
criterion to select daytime observations out of twilight measurements is that
we require the tangent point solar zenith angle to be less than or equal to
88∘. Using this approach we avoid twilight measurements and all of
the measurement points are located on the dayside.
Examples of the daytime-calibrated spectra for orbit number 41 455, measured
on 3 February 2010 at a mean latitude of 17.3∘ N and a mean
longitude of 94.3∘ E, are shown in Fig. a for the
O2(1Σ) band and in Fig. b for the
O2(1Δ) band.
The spectral region used to observe the daytime O2(1Σ) spectrum
includes a Rayleigh-scattering background which perturbs the retrieval.
Consequently it is necessary to estimate the Rayleigh scattering and subtract
it from the observational spectrum to yield the O2(1Σ) emission spectra. This
background scattering is attributed in part to the upwelling radiation,
multiple scattering in the lower atmosphere, and the terrestrial albedo. This
results in an absorption signature for O2(1Σ)
. To account for the multiple scattering and absorption
from the ground state O2(3Σg) to the
O2(1Σ), a background signal comprising the O2(1Σ)
spectrum at the highest altitude (≈ 148 km) scaled to the ratio of
the mean of the out-of-band radiances is subtracted from the limb spectra at
each tangent height. We consider the spectra in the 750–759 and 767–780 nm
as out of band.
After this correction, we subtract a linear background from the whole signal
in each level. An example of the background-subtracted spectrum containing
the O2(1Σ) emission is shown in Fig. c.
Examples of the daytime-calibrated spectra and the
background-corrected spectra. (a) O2(1Σ) on
3 February 2010; orbit 41 455; mean latitude 17.3∘ N, mean
longitude 94.3∘ E. (b) as (a) but for
the O2(1Δ) band. (c) as (a) but with background
correction applied. (d) as (b) but with background
correction applied.
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For the O2(1Δ) band, the absorption signature in the spectral
background is negligible compared to the daytime O2(1Σ) band
spectra; therefore, we only subtract a linear background from the observation.
An example for the daytime O2(1Δ) band spectra with background
subtracted is shown in Fig. d.
Twilight spectra
Because the tangent point solar zenith angle for the sun being below the
horizon varies with tangent altitude, we use Eq. () to calculate
the horizon angle for each tangent point (R is the radius of the Earth and
h is the tangent point height):
αhorizon=π2+cos-1RR+h.
As a criterion to select twilight data, we remove every limb scan in which at
least one point measurement has a solar zenith angle less than given by
Eq. (). Based on this, we obtain the O2(1Σ) and the
O2(1Σ) background-subtracted twilight spectra shown in
Fig. a and c for the example orbit (41 455). For the twilight
O2(1Δ) band, we apply the same background subtraction as for
daylight. Figure b shows the twilight O2(1Δ)
spectra, and panel (d) shows the background-corrected twilight
O2(1Δ) spectra for the same example orbit (41 455). It is
apparent that the background signal is negligible for both of the
O2(1Σ) and O2(1Δ) twilight spectra.
Examples of the twilight-calibrated spectra and the spectra from
which the background O2(3Σg) to O2(1Σ)
absorption has been subtracted. (a) for the O2(1Σ) band on 3 February 2010; orbit
number 41 455; mean latitude of 78.0∘ N, mean longitude of
226.5∘ E.
(b) as (a) but for the O2(1Δ) band.
(c) as (a) but with background subtracted.
(d) as (b) but with background subtracted.
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Retrieval
To invert the observed radiation to spectral emission rates, we set up
30 layers around the Earth, such that each layer is centered at one tangent
height. We denote the observed spectral radiances by Y, the path
length of each of the observed lines of sight through each of the atmospheric
layers by L, and the emission rate from the layer by X.
Assuming no self-absorption, this yields the linear relation
Y=LX.
In Eq. (), dim(Y) = number of atmospheric
layers × number of spectral points = 30×144 for the
O2(1Σ) band and 30×210 for the O2(1Δ)
band. Dim(L) = number of the atmospheric
layers × number of tangent heights in each scan = 30×30. Dim(X) = number of atmospheric layers × number
of spectral points in the corresponding wavelength interval = 30×144 for O2(1Σ) and 30×210 for O2(1Δ).
We solve Eq. () by minimizing
||Y-LX||2 using standard least squares and
normalization with the error covariance matrix Sy and obtain the
inverted spectral emission intensity X:
X=(LTSy-1L)-1LTSy-1Y.
The error covariance matrix Sy has diagonal elements of the
out-of-band variances of the background-corrected spectra in each altitude.
Results
Inverted spectra
Using our method in Sect. , the emission intensities
are calculated for the spectra described in
Sect. and . Examples of
the emission intensity for daytime O2(1Σ) are shown in
Fig. a and for daytime O2(1Δ) in
Fig. b.
The spectral shapes of the O2(1Σ) band in the daytime
O2(1Σ) spectra and of the O2(1Δ) band in the
daytime O2(1Δ) spectra are clearly visible. We find that the
largest values for the daytime O2(1Σ) band are located at about
90 km altitude, and for the O2(1Δ) band at 54 km, noting that
the SCIAMACHY MLT scan range is 50 to 150 km. We also note that the maximum
O2(1Δ) emission intensities in each limb scan are about 2 orders of magnitude larger than the maximum values of the
O2(1Σ) band in the corresponding limb scan.
Figure c shows the twilight O2(1Σ) emission
intensity, and Fig. d shows the twilight O2(1Δ)
emission intensity for the same orbit but retrieved from one of the three
twilight MLT scans (see Sect. ). The
O2(1Σ) band is about 1 order of magnitude smaller than during
daylight. The twilight O2(1Δ) band signal is more pronounced
but about 2 orders of magnitude smaller than during the day. The error bars
in the panels of Fig. represent the square root of each of the
diagonal elements of the retrieval error covariance matrix
Sa for each altitude:
Sa=GSyGT,
in which the contribution function matrix G is defined as
G=LTSy-1L-1LTSy-1.
Evaluating all altitudes, not shown here but indicated in
Fig. a, we observe the strongest signal of daytime
O2(1Σ) around 83–99 km. Hereafter, we use the term
“significant” for data with a signal-to-noise ratio (SNR) greater than 1.
Figure c shows that the twilight O2(1Σ) band
emission intensities have large noise masking the signal. The daytime
O2(1Δ) emission intensities are in general of higher SNR than
the twilight O2(1Δ) emission intensities. The daytime
O2(1Δ) emission intensities are strongest at the lowest
observable altitudes, i.e., 54 km (Fig. b). The strongest
twilight O2(1Δ) emissions are located in the 83–96 km
altitude range (Fig. d shows only a selection of altitudes).
Examples of the emission intensities that are obtained by solving
Eq. (). (a) For the daytime
O2(1Σ) band on the date 3 February 2010; orbit number
41 455; mean latitude of 17.3∘ N, mean longitude of
94.3∘ E. (b) as (a) but for the O2(1Δ)
band. (c) For the twilight O2(1Σ) band on the date
3 February 2010; orbit number 41 455; mean latitude 78.0∘ N, mean
longitude 226.5∘ E. (d) as (c) but for
the O2(1Δ) band. Error bars represent the retrieval errors.
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Typical profiles of the VER for different latitudes.
(a) For the daytime O2(1Σ) VER on the date
3 February 2010; orbit number 41 455. (b) as (a) for
the daytime O2(1Δ) band. Error bars represent the retrieval
errors.
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Volume emission rates
We integrate the spectral emission intensity from 759 to 767 nm to obtain
the O2(1Σ) band integrated volume emission rate. Volume
emission rate profiles for one sample satellite orbit (41 455 on
3 February 2010) for daytime O2(1Σ) and O2(1Δ)
are shown in Fig. a and b respectively. Examples of the volume
emission rate latitude–altitude distributions for the same orbit for daytime
O2(1Σ) are shown in Fig. a and for daytime
O2(1Δ) in Fig. b. The blank regions represent
areas with signal-to-noise ratios of less than 1.
Latitude–altitude contours of the daytime VER. (a) for the
O2(1Σ) band and for one satellite orbit on the date 3
February 2010; orbit number 41 455. (b) as (a) for
O2(1Δ). Signal-to-noise ratios less than 1 and large noisy
values are excluded. (c) and (d) as
(a) and (b) respectively, averaged on all of the orbits on
the whole day of 3 February 2010, with the same logarithmic scale. Negative
values are excluded. (e) Signal-to-noise ratios of the daily mean of
O2(1Σ) VERs. Areas where the signal-to-noise ratio is less
than 1 are plotted in white. (f) as (e) for the
O2(1Δ) band.
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The O2(1Σ) VER has its maximum in the 90–98 km altitude
range, which is 2 orders of magnitude smaller than the O2(1Δ)
maximum VER, as shown in Fig. a. The volume emission rate
profile of dayglow O2(1Δ) observed by TIMED–SABER often has
its maximum around 50 km altitude, as shown for example in Fig. 1 of
. Figure b shows that the SCIAMACHY
MLT volume emission rate profiles are largest at the bottom of the observed
altitude range, around 54 km. These VER profiles sometimes show secondary
maxima in the range 80–90 km, which are at least 1 order of magnitude
smaller than the largest SCIAMACHY VER. This secondary maximum occurs
especially around equinox times. The measurement errors of the volume
emission rates (not shown here) for different orbits show that the
O2(1Σ) volume emission rates are significant from 65 to
140 km and do not depend on latitude. Inspecting the volume emission rate
altitude profiles (not shown here), we see that the O2(1Σ)
volume emission rates have the largest SNR below 125 km and above 85 km.The
best signal of the O2(1Δ) VER is below 95 km.
In a region above the South Atlantic and off the Brazilian coast, the Earth's
magnetic field is anomalously low and the ionizing radiation can be increased
by several orders of magnitude. This region is called the South Atlantic
Anomaly (SAA, see for example ) and any spacecraft
which crosses this region can give false instrument readings. In our
retrievals, the SNRs of the volume emission rates in the orbits that cross
this region are affected by the SAA. The most dramatic influence of the SAA
on our dataset is on the O2(1Σ) volume emission rates SNR,
although the values are still significant in the 80–100 km altitude range.
The SAA influences the O2(1Σ) measurements more than the
O2(1Δ) measurements.
Time series of the daily mean VER. (a) for the daytime
O2(1Σ) VER; 30∘ N; July 2008 to March 2012.
(b) as (a) for the daytime O2(1Δ) band.
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Daily mean VER latitude–altitude distributions
We calculate the daily mean VERs as follows. We bin the measurements into
5∘ latitude bins. In each bin, the measurements located within
±2.5∘ are attributed to that latitude and averaged to the daily mean
VER. An example of the daily mean daytime O2(1Σ) VER
latitude–altitude distribution (on 3 February 2010) is shown in
Fig. c and of the daily mean daytime O2(1Δ) VER
in Fig. d (VERs with low signal-to-noise ratios and with large
measurement errors are excluded). The daily mean O2(1Σ) VERs
have maximum values of about 1–2 orders of magnitude smaller than
O2(1Δ). Similar to our results for a single orbit
(Sect. ), we observe the largest
O2(1Δ) VER below 60 km and the largest O2(1Σ)
VER at 90 km.
To assess the signal-to-noise ratio for the daytime VERs,
Fig. e shows the daily mean O2(1Σ) VER signal-to-noise ratios. We observe the strongest signal of daytime O2(1Σ)
in the 70–130 km altitude range. The strongest signal of the twilight
O2(1Σ) is observed between 84 and 95 km (not shown here).
Figure f shows that the stronger signal of daytime
O2(1Δ) is observed below 105 km, with the strongest around
70 km. The largest signal of twilight O2(1Δ) is observed in
the altitude range of 83–97 km (not shown).
Time series
In the following, we discuss the variation of the daily mean VERs versus
latitude and time. First we will discuss the temporal variation in the
mesosphere and lower thermosphere (70–140 km) at 30∘ N. Thereafter
we will discuss the variation of the peak values, peak altitudes, and
centroid altitudes as a function of time and latitude.
Time series at 30<inline-formula><mml:math id="M282" display="inline"><mml:msup><mml:mi/><mml:mo>∘</mml:mo></mml:msup></mml:math></inline-formula> N
By calculating the daily mean VERs for all of the days on which SCIAMACHY MLT
limb scans are available, we obtain time series of the daily mean VERs from
July 2008 to March 2012. An example of these time series of the daily mean
daytime O2(1Σ) VER, chosen for 30∘ N from all
altitude combinations, is shown in Fig. a and for the daytime
O2(1Δ) VER of the same location in Fig. b.
We found a semi-annual variation with the strongest O2(1Σ)
signal in the 90–95 km range during May–June and September–November and
the lowest signal in December–March, with a secondary minimum in August. The
highest values of the O2(1Δ) VER are located at the lowest
altitude of observations, formed mostly by ozone photodissociation
. We observe secondary maximum values which mostly
occur in May–June and September–November (approximately spring and autumn)
in the 75–95 km altitude range. The secondary maximum of
O2(1Δ) occurs in the same altitude range and with the same
temporal variation as the O2(1Σ) signal. This will be
investigated in more detail in the following section.
Variation of peak values
Next we evaluate the variations of the maximal daily mean VERs in the
mesosphere and lower thermosphere with respect to latitude and time. For
this, we derive the maximal values from the daily mean VERs for
O2(1Σ) and between 85 and 100 km altitude for
O2(1Δ), which are shown in Fig. a and b
respectively. Only those regions are shown in Fig. b, d, and f
where the secondary maxima of O2(1Δ) exist. The peak altitudes
are also obtained at the same time and will be discussed in
Sect. . The maxima of the
O2(1Σ) VER at middle to low latitudes
(60∘ S–60∘ N) appear to be correlated with the maximum
intensity of solar radiance. Additionally, we sometimes observe attenuations
in the maximum values in the late northern winters, mostly from late January
until early March of each year. Also there are some high values at northern
polar latitudes: in spring 2009, from autumn to spring 2009–2010, from
autumn to winter 2010, and from autumn to spring 2011–2012.
The secondary maxima of the O2(1Δ) VER are confined to winter
at mid-to-high latitudes.
The correlation of the O2(1Σ) VERs with the solar radiance
suggests formation caused by solar light, either by ozone photolysis
(Reactions and ) or by larger
abundances of atomic oxygen owing to stronger O2 photolysis. In the second
case, the formation of the excited states could also be due to the
recombination of atomic oxygen
(Reactions and ). The maximal
values at high latitudes, where the solar flux is low, in general suggest
other sources such as recombination of atomic oxygen .
This is true in particular for O2(1Σ), where high values occur
at high latitudes in winter and spring; these are probably caused by downward
transport of thermospheric atomic oxygen into the mesosphere and lower
thermosphere.
There are two possibilities for secondary maxima of O2(1Δ).
They happen in the region where the secondary ozone maximum is strongest.
Also, atomic oxygen densities might be strongest due to enhanced mixing with
the lower thermosphere. Detailed study of the processes which result in the
formation of the secondary maxima of O2(1Δ) is beyond our work.
Variation of peak altitudes
The altitudes of the peak values of O2(1Σ) and
O2(1Δ) are shown in Fig. c and d. The altitudes
of the peak values of O2(1Σ) roughly follow the maximum
intensity of solar radiance but show highest values at low-to-middle
latitudes. We refer to the maximum intensity of solar radiance as the
latitudes and times in which solar zenith angles have their lowest values.
The low values of the maximum VERs and the high altitudes of the maximal VERs
at the outermost high latitudes in the northern winters correspond to low
signal-to-noise ratios and are below our significance level.
In the regions where the secondary maxima of O2(1Δ) happen, the
peak altitudes occur in the ∼ 84–89 km altitude range.
(a) Time series of the maximal daily mean
O2(1Σ) VER. (b) as (a) for
the O2(1Δ) band. (c) Time series of the altitudes of the
maximum daily mean O2(1Σ) VER. (d) as (c)
for the O2(1Δ) band. (e) Time series of the centroid
altitudes of the daytime O2(1Σ) daily mean VER (km).
(f) as (e) for the O2(1Δ) band.
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Centroid altitudes
The estimation of the peak altitude is affected by instrument noise and
vertical resolution. A more stable measure is the centroid altitude
hCA:
hCA=∑iaivi∑ivi,
where vi is the volume emission rate at altitude ai, which ranges
from ∼ 50 to ∼ 150 km for O2(1Σ) and from
∼ 85 to ∼ 100 km for O2(1Δ).
Figure e shows the centroid altitude for daytime
O2(1Σ) and Fig. f for daytime
O2(1Δ).
The maximum hCA values of the O2(1Σ) band are correlated
with the (maximum intensity) of solar radiance. We also observe maximal
values of hCA in a narrow band at northern polar latitudes, where
the solar flux generally is low. This also suggests recombination of atomic
oxygen as a source, as discussed in the Sect. .
The very low values of the hCAs, at the highest latitudes, mostly
in the hemispheric wintertimes, correspond to low signal-to-noise ratios of
the corresponding VERs (not shown here) and are below our significance level.
The O2(1Δ) secondary maximums occur in winter at high
latitudes. The values of the O2(1Δ) hCA range
between ∼ 88 and ∼ 89 km altitude.
Figure c, d, and f show a decrease in the altitude of the
maximum O2(1Σ), altitude of the maximum O2(1Δ),
and O2(1Δ) hCA respectively between November 2010
and February 2011. This is due to a change in the limb sequence so that
tangent altitudes were shifted, as seen for example at lower altitudes in
these days. However, this does not have any notable effect on the VER time
series in Figs. a, b, a, and b.
Discussion of temporal–spatial variation
The temporal–latitudinal variation in peak values and altitudes suggests that
O2(1Σ) at 85–100 km altitude is formed by a combination of
ozone photolysis and atomic oxygen recombination. At high latitudes during
winter and spring, atomic oxygen recombination dominates, but, in the
subsolar region, ozone photolysis is more important. In contrast, the
secondary peak of O2(1Δ) stems mainly from atomic oxygen
recombination, in particular at high latitudes during winter.
Modeled VERs of O2(1Σ) (a) and
O2(1Δ) (b). Solid lines are the mean of all model
results for solar zenith angles 0, 10, 20, 30, 40, 50, 60, 70, 80, and
88∘ considering all formation processes; dotted lines show the
variability due to photolysis reactions (± standard deviation). The
dashed, dash-dotted, and dash–dot–dot lines show the contributions of
individual formation processes: O+O, O(1D) quenching, and
resonant excitation for O2(1Σ), O+O,
O2(1Σ), and ozone photolysis for O2(1Δ). Black
lines with symbols are SCIAMACHY profiles at five latitudes observed on
3 February 2010 (see Fig. ).
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The temporal–latitudinal variation in peak values and altitudes suggests that
the daytime VERs of O2(1Δ) and O2(1Σ) in the
altitude range of 80–100 km are formed by a combination of ozone and
O2 photolysis and atomic oxygen recombination. At high latitudes
during winter and spring, atomic oxygen recombination dominates, but, in the
subpolar region, photolysis of ozone and O2 is more important.
To test whether these conclusions are generally consistent with our
understanding of the photochemical production and loss of
O2(1Δ) and O2(1Σ), a simple photochemical model
covering the production and loss reactions summarized in
Sect.
(Reactions –)
was set up. Photochemical equilibrium was considered for
O2(c1Σu-), O(1S), O(1D),
O2(1Δ), and O2(1Σ). Reaction rates were taken
from the JPL recommendation with the exception
of the quenching of the intermediate O2(c1Σu-)
state (the q6 and q7 rates shown in Fig. ), which was
taken from with δ coefficients from
. Einstein coefficients for the
O2(1Σ) → O2(X3Σg-)
transition (A2 in Fig. ) were taken from
, for the
O(1S) → O(1D) transition (A4 in
Fig. ) from the NIST atomic spectra database
, for the
O(1S) → O(3P) transition (A5 in
Fig. ) from the NIST atomic spectra database, and for the
O2* products (not given in Fig. ) from
. Photolysis rates for solar zenith angles 0, 10,
20, 30, 40, 50, 60, 70, 80, and 88∘ were calculated using a fixed
ozone profile using the 3dCTM model . The rate
of photoexcitation of the ground state O2(X3Σg-)
to the second excited state O2(1Σ) was calculated following
but using recent line-strength data provided by the
HITRAN database at http://hitran.org ,
yielding a rate of 2.04×10-9 photons cm-3 s-1 above
70 km. Temperature, total air density, and the densities of O2 and
N2 were taken from the NRLMSISE-00 model
at 10:00 local time at the equator. The production of O(1D) by
photolysis of H2O and CO2 in the Ly-α range was also
considered, and H2O and CO2 were assumed to have constant
mixing ratios of 1 ppm (H2O) and 380 ppm (CO2). Ozone
density was adapted from a multi-year global average of SABER data
in the altitude region 70–104 km (see ,
Fig. 3). Atomic oxygen was calculated from photochemical equilibrium of ozone
considering only ozone photolysis and production by O+O2. The
resulting VERs of O2(1Δ) and O2(1Σ) are shown in
Fig. , compared to the SCIAMACHY daytime data of
3 February 2010 (shown in Fig. ). Considering that the ozone
profile, O2, N2, and temperature are not chosen to fit those
specific observations, the agreement is very good for both emissions,
indicating that the main processes of O2(1Δ) and
O2(1Σ) formation and loss are reproduced well by this simple
model. Also shown are the contributions of the individual production
reactions: O + O, quenching of O(1D), and resonance excitation of
O2(X3Σg-) for O2(1Σ), O+O,
quenching of O(1Σ), and ozone photolysis for
O2(1Δ).
Below about 82 km, O2(1Σ) is formed in about equal amounts by
quenching of O(1D), while above quenching of O(1D)
dominates. The reaction of O+O contributes about 1 order of
magnitude less O2(1Σ) than the other two branches even in the
region where it has the largest contribution (around 90 km). This is
consistent with the ratio of emission intensities during twilight and during
daytime of about a factor of 10 as discussed in
Sect. , considering that, during nighttime and
twilight, O2(1Σ) is formed solely by the O + O reaction.
The formation of O2(1Δ) is dominated by ozone photolysis at all
altitudes, though below 90 km O2(1Σ) quenching contributes
about 10–25 %.
O(1D) is formed mainly by photolysis of O3 below 90 km and
by photolysis of O2 above that altitude. During daytime, both
the O2(1Δ) and O2(1Σ) states
are formed by photolysis in agreement with
our observation that the peak maxima correlate with the maximum intensity of
solar radiance. The contribution of the O+O reaction generally is
smaller by 1 to 3 orders of magnitude than the contribution of photolysis, in
agreement with the lower VERs observed during twilight when photolysis is not
a significant production process. However, it should be pointed out that the
relation between the O2 airglow and atomic oxygen by the O+O
reaction and de-excitation of the excited intermediate states such as
O2(c1Σu-) is probably quadratic to cubic;
increased amounts of atomic oxygen, e.g., at high latitudes during winter
when large amounts of atomic oxygen can be transported or mixed down from the
lower thermosphere, can therefore increase the contribution of the O+O
reaction to the overall airglow considerably. This explains the observed
enhancements of the airglow at high northern latitudes in winter and spring,
in particular as the centroid altitudes in these occasions range around
90 km, the altitude where the O+O reaction has the strongest
influence.
(a) Daily mean of the daytime O2(1Σ) VER
from December 2008 to April 2009 averaged between 60 and 70∘ N
latitude. (b) as (a) for the O2(1Δ) band.
(c) Daily mean of the signal-to-noise ratios of the daytime
O2(1Σ) VER from December 2008 to April 2009 averaged between
60 and 70∘ N latitude. (d) as (c) for
O2(1Δ) VER signal-to-noise ratios.
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Relationship between the VER time series variations and sudden stratospheric warmings – example
Sudden stratospheric warmings are dynamical phenomena in the winter
polar stratosphere caused by upward propagating planetary waves interacting
with the mean flow . During the so-called recovery phase
of SSWs, the reformation of the jet changes gravity wave propagation to the
mesosphere.The induced change in the residual circulation results in an
enhanced descent of air. This causes adiabatic warming and the stratopause
reforms at altitudes as high as 75–80 km . The unusual
brightening of the OH airglow is presumably caused
by enhanced downwelling of atomic oxygen. According to , a
major SSW event happened around 21 January 2009 (Fig. 1 of that paper). We
therefore expect to observe enhanced airglow at the end of January 2009.
Based on the temporal evolution of mesospheric temperature during the SSW
event, divided the response in the mesosphere into
three stages: the period prior to day 15 is considered the normal stage, days 15–22
correspond to the cooling stage, and days post-22 correspond to the recovery
stage. According to this, they reported that the O2 nightglow
brightness decreased by about a factor of 10 during the cooling stage and
then increased by about a factor of 3 during the recovery stage relative to
the normal stage. Figure a and b show the time series from
December 2008 to April 2009 of the daytime O2(1Σ) and
O2(1Δ) daily mean VER, averaged from 60 to 70∘ N. We
observe that on the last day of the cooling stage, i.e., 22 January 2009, the
daily zonal mean O2(1Σ) and O2(1Δ) VERs show a
reduced maximum intensity in the 82–87 km altitude range. We observe larger
intensities on the next measurement day of SCIAMACHY about 3 weeks later,
i.e., on 10 February 2009. This is expected from a decrease of atomic oxygen
due to horizontal mixing and upwelling during the cooling stage and then
downward extension of the MLT region with large mixing ratio of O
during the recovery stage of the SSW . On the
measurement day of SCIAMACHY after the recovery phase, i.e., 23 March 2009,
the relative difference in the O2(1Δ) signal is less prominent
compared to the relative difference in O2(1Σ). A detailed
analysis of this relationship is beyond the scope of this paper.
We also note that the signal-to-noise ratios of the VERs for
O2(1Σ) shown in Fig. c are statistically
significant. It is shown in Fig. d that, during and after the
initial stage of the SSW event, the O2(1Δ) signal becomes
weaker and stronger respectively, and the signal-to-noise ratio of the data
are such that this behavior is statistically significant.
Discussion and conclusions
We present the retrieval of daytime and twilight O2(1Σ) and
O2(1Δ) spectral emissions from MLT measurements of the airglow
in limb-viewing geometry from the SCIAMACHY instrument onboard Envisat. From
the retrieved spectra, we calculate the band integrated VERs for both bands.
The maxima of the O2(1Σ) VER and the centroid altitude of the
O2(1Σ) are correlated with the maximum intensity of solar
radiance. High values of maximum VER and centroid altitude are additionally
seen at northern polar latitudes. The (30∘ N) time series of
O2(1Σ) VER shows a maximum in the 90–98 km altitude range.
The maximum values correspond to high centroid altitudes for
O2(1Σ). The daily zonal (60–70∘ N) mean
O2(1Σ) VERs show a reduced maximum intensity in the 82–87 km
range and in the initiation of the sudden stratospheric warming 2009 event as
well as an increase in intensity about 3 weeks later.
The maxima of the O2(1Δ) VERs are also correlated with the
maximum intensity of solar radiance. They are most prominent in summer, while
the altitudes of the maximal values have their highest values in winter, and
both occur at high latitudes. The time series of O2(1Δ) VER is
2 orders of magnitude larger than O2(1Σ) VER at its maximal
values which are located below the observation altitude (< 60 km), but it
shows some secondary maxima about 1 order of magnitude smaller than the
primary maxima at 84–89 km. This happens in winter at high latitudes. The
maximum VERs correspond to the low centroid altitudes for the
O2(1Δ) band. The daily zonal (60–70∘ N) mean
O2(1Δ) VERs show a reduced maximum intensity in the 82–87 km
range and in the initiation of the sudden stratospheric warming 2009 event as
well as an increase in intensity about 3 weeks later, although the relative
difference in the O2(1Δ) band signal is less prominent compared
to the relative difference in O2(1Σ).
The intensification of the VER during the sudden stratospheric warming in
early 2009 presumably corresponds to the downwelling of the atomic oxygen
following the warming, while the decrease is probably due to upwelling as
well as horizontal mixing during the warming event.
Our results suggest that at low and middle latitudes O2(1Σ) and
O2(1Δ) abundances during daytime are dominated by photolysis of
ozone below 90 km and by photolysis of ozone and O2 above 90 km as
supported by our observed correlation with solar illumination, and they are consistent
with the processes depicted in Fig. . At high latitudes,
however, in particular during winter, atomic oxygen abundances might be a
more important driver due to the recombination of O+O and subsequent
de-excitation via O(1S) and O2(1Σ).
As the formation of the O2(1Δ) band is dominated by ozone
photolysis, this band can be used to derive ozone densities directly.
However, as quenching of O2(1Σ) contributes about 10–25 %
to the overall production in 70–90 km, the accuracy of this retrieval can
be improved considerably when both lines are available. The
O2(1Σ) band is dominated by quenching of O(1D), and
daytime O(1D) can be derived from observations of the
O2(1Σ) VER. However, below ∼ 90 km, the resonant
excitation from the ground state has to be taken into account as well. The
rate used here based on new line-strength data from HITRAN is lower by about
a factor of 2 compared to a similar estimate used for HRDI data as
described in and lower by about a factor of 4
compared to the original estimate by . We conclude
that the retrieval of O(1D) and ozone as the main source of
O(1D) below 90 km from O2(1Σ) is possible but needs a
careful estimation of the rate of resonant excitation.