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Measurements of stratospheric volcanic aerosol optical
depth from NOAA TIROS Observational Vertical
Sounder (TOVS) observations
Clémence Pierangelo, Alain Chédin, Patrick Chazette
To cite this version:
Clémence Pierangelo, Alain Chédin, Patrick Chazette. Measurements of stratospheric volcanic aerosol optical depth from NOAA TIROS Observational Vertical Sounder (TOVS) observations. Journal of Geophysical Research: Atmospheres, American Geophysical Union, 2004, 109 (D3), pp.n/a-n/a. �10.1029/2003JD003870�. �hal-02902720�
Measurements of stratospheric volcanic aerosol optical depth from
NOAA TIROS Observational Vertical Sounder (TOVS) observations
Cle´mence Pierangelo and Alain Che´din
Laboratoire de Me´te´orologie Dynamique, Institut Pierre Simon Laplace, Ecole Polytechnique, Palaiseau, France
Patrick Chazette
Laboratoire des Sciences du Climat et de l’Environnement, Institut Pierre Simon Laplace, Commissariat a` l’Energie Atomique, Gif-sur-Yvette, France
Received 17 June 2003; revised 18 September 2003; accepted 24 November 2003; published 14 February 2004.
[1] We show that the infrared optical depth of stratospheric volcanic aerosols produced by the eruption of Mount Pinatubo in June 1991 may be retrieved from the observations of the High-Resolution Infrared Radiation Sounder (HIRS-2) on board the polar meteorological satellites of the National Oceanic and Atmospheric Administration (NOAA). Evolution of the concentration in time and in space, in particular the migration of the aerosols from the tropics to the Northern and Southern Hemispheres, is found to be consistent with our knowledge of the consequences of this eruption. The method relies on the analysis of the differences between the satellite observations and simulations from an aerosol-free radiative transfer model using collocated radiosonde data as the prime input. Thus aerosol optical depths are retrieved directly without making assumptions about the aerosol size distribution or absorption coefficient. The aerosol optical depths reached a maximum in August 1991 in the tropical zone (0.055 at 8.3mm, 0.03 at 4.0 mm, and 0.02 at 11.1mm). The peak occurred in November 1991 in the southern midlatitudes and in March/April 1992 in the northern midlatitudes. A reanalysis of the almost 25 year archive of NOAA TIROS-N Operational Vertical Sounder (TOVS) observations holds
considerable promise for improved knowledge of the atmosphere loading in volcanic aerosols. INDEXTERMS: 0305 Atmospheric Composition and Structure: Aerosols and particles (0345, 4801); 0370 Atmospheric Composition and Structure: Volcanic effects (8409); 3360 Meteorology and Atmospheric Dynamics: Remote sensing; KEYWORDS: volcanic aerosols, satellite remote sensing, infrared optical depth
Citation: Pierangelo, C., A. Che´din, and P. Chazette (2004), Measurements of stratospheric volcanic aerosol optical depth from NOAA TIROS Observational Vertical Sounder (TOVS) observations, J. Geophys. Res., 109, D03207, doi:10.1029/2003JD003870.
1. Introduction
[2] Volcanic eruptions, by producing significant
transito-ry solar and infrared radiative perturbations, are recognized as an important natural cause of climate variability [Stenchikov et al., 1998]. Identifying and quantifying natu-ral fluctuations is essential to separate them from anthropo-genic fluctuations [Robock, 2000; Ramachandran et al., 2000; Kirchner et al., 1999; Dutton and Christy, 1992].
[3] In June 1991, the eruption of Mount Pinatubo
(Philip-pines) injected about 20 Mt of sulfur dioxide, primarily into the stratosphere. The rapid conversion of sulfur dioxide into sulphuric acid (H2SO4) droplets caused a significant
extinc-tion, both in the visible and in the infrared. These particles were steadily removed and about three years after the eruption all the aerosols were gone. An important consequence of the eruption was a warming of the stratosphere (by about 2 – 3 K) and a cooling of the troposphere (by about 0.5 K).
This eruption has been the object of many studies using in situ or satellite observations [McCormick et al., 1995], both in the visible/ultraviolet and infrared spectral domains.
[4] Observations of the atmosphere in the infrared
spec-tral region have several advantages. First, during the night or over high-albedo surfaces [Ackerman, 1997], visible techniques cannot be applied, contrary to infrared tech-niques. Second, with high-spectral-resolution sounders, like the Aqua/Advanced Infrared Sounder (AIRS) or the Metop/ Infrared Atmospheric Sounder Interferometer (IASI), it should be possible to identify spectral signatures of the aerosol components. Third, with instruments like the NOAA TIROS-N Operational Vertical Sounder (TOVS), first launched in 1978 on board the NOAA polar meteoro-logical satellite series, an archive of 25 years of observa-tions is now available. Moreover, global infrared aerosol optical depths are required to compute accurate radiative fluxes and budget. As few direct measurements of aerosol thermal infrared optical depths are available, the general approach is to compute them from aerosol size distribution and refractive indices with a Mie code [Dutton, 1995].
Copyright 2004 by the American Geophysical Union. 0148-0227/04/2003JD003870
[5] Lambert et al. [1993] studied Pinatubo aerosol using
satellite observations in the infrared from the Improved Stratospheric and Mesospheric Sounder (ISAMS). They retrieved area weighted global mean stratospheric optical thickness at 12.1 mm from the zonal mean extinction profiles for the period going from November 1991 to April 1992. The Cryogenic Limb Array Etalon Spectrometer (CLAES) instrument made measurements of the infrared emission of stratospheric aerosol near 12mm from October 1991 until May 1993. Mergenthaler et al. [1995] obtained the aerosol distribution of absorption cross section, optical depth and sulfuric acid mass estimates with CLAES mea-surements. Lambert et al. [1997] combined CLAES and ISAMS to retrieve aerosol composition, volume and area mass density. Baran et al. [1993] analyzed the effect of Pinatubo aerosols on differences in brightness temperatures between High-Resolution Infrared Radiation Sounder (HIRS-2) channels of two NOAA satellites (channel 12.5 mm of NOAA 11 and channel 8.3 mm of NOAA 12) using in situ observations of aerosol particle size and number density. They suggested to retrieve the aerosol mass loading from this brightness temperature difference. Ackerman and Strabala [1994] used three brightness tem-peratures (at 8, 11, and 12mm from HIRS-2), and, assuming a chemical composition and a size distribution of the aerosol from in situ observations, retrieved their visible and infrared optical depths. Halogen Occultation Experiment (HALOE) instruments have been used to retrieve aerosol extinction profile using solar occultation techniques [Hervig et al., 1995]. This method obviously applies to near infrared (from 2.45 mm to 5.26 mm). Echle et al. [1998] also determined optical and microphysical parameters of Pinatubo aerosols using the Michelson Interferometer for Passive Atmospheric Sounding, Balloon-borne version (MIPAS-B). Although this midinfrared limb sounder was not yet on board a satellite platform, at that time, and provided data for only two days (14 – 15 March 1992), it showed the potential of using infrared data to complement measurements in the visible spectrum and retrieve chemical composition and size distribution of volcanic aerosols.
[6] The TOVS instrument consists of three passive
verti-cal sounder radiometers [Smith et al., 1979]: HIRS-2 with 19 channels in the infrared band and one in the visible band, the Microwave Sounding Unit (MSU), a microwave radiometer with four channels in the vicinity of 55 GHz, and a pressure-modulated infrared radiometer, the Stratospheric Sounding Unit (SSU), with three channels near 15mm. Only HIRS-2 data are used here. Scan widths are approximately 2200 km wide, providing global coverage every 12 hours. HIRS-2 measures atmospheric and/or surface emission in seven channels located around 15 mm, five located around 4.3 mm, one window channel at 11 mm, and three water vapor channels located around 6.7 mm. Surface and ozone emission is measured in one 9.6 mm channel, surface emission and reflected solar radiation in two 3.7mm chan-nels, and reflected solar radiation in one visible channel.
[7] In this paper, we first show that HIRS-2, on the two
platforms NOAA 10 (from July 1987 to September 1991) and NOAA 12 (from July 1991 to August 1995), is capable of identifying the signature of the Mount Pinatubo aerosols. These signals can clearly be revealed by a careful analysis of the difference between radiance observations from
HIRS-2 and simulations from an aerosol-free radiative transfer model using space-time collocated radiosonde tem-perature and humidity measurements as the prime input. These experimental signatures are then compared with theoretical signatures obtained from an aerosol-loaded radi-ative transfer model. This procedure is applied to four channels chosen for their different spectral sensitivities to volcanic aerosols (channel 18 at 4.0 mm, channel 10 at 8.3mm, channel 8 at 11.1 mm, and channel 5 at 14 mm). In a second step, we interpret theses signatures in terms of aerosol optical depths with no assumption on the aerosols properties and analyze their time and space variations. The optical depths retrieved at three wavelengths (11.1 mm, 8.3 mm and 4 mm), and for three zones of latitude (20 – 60N, 20S to 20N, 60 – 20S), are shown to be in the range of previous estimations. Given the size distribution of the Pinatubo aerosols from in situ measurements by Deshler et al. [1993] and H2SO4 refractive index [Biemann et al.,
2000], the particle number concentration is then retrieved. The time variation of the optical depth clearly reflects the decay and migration of the aerosols from the tropics to the Northern and Southern Hemispheres.
2. Signature of Mount Pinatubo Eruption Aerosols
[8] Figure 1 shows the absorption cross section of
Pina-tubo aerosols at the 19 HIRS-2 wavelengths in the infrared from 3.70 to 15.0 mm. The absorption cross section is calculated with a Mie code, assuming the refractive indices of Biemann et al. [2000] (215 K, 75% H2SO4) and two size
distributions. The first one is a lognormal bimodal distribu-tion with mode radii 0.2mm and 0.6 mm, mode width 1.6 and 1.3 and their relative abundance 0.95 and 0.05. Balloon-borne measurements show that mode radii and width vary with altitude and time [see Deshler et al., 1993]. Neverthe-less, the previous values, which are temporal and spatial averages of in situ measurements are chosen as being the most representative of the whole period studied here. The Figure 1. Absorption cross section computed with Mie theory for the 19 HIRS-2 infrared wavelengths, indicated by crosses, for two size distributions of volcanic aerosols (monomodal from Chazette et al. [1995] and bimodal from Deshler et al. [1993]).
second distribution is the equivalent lognormal monomodal distribution with cross radius 0.45 mm and width 1.7 mm, previously determined by Chazette et al. [1995] from Deshler et al. [1993] in situ measurements. As expected, results are similar, with a maximum of the absorption cross section for channel 10, at 8.3 mm, corresponding to the strongest absorption band of H2SO4.
2.1. Collocated Radiosonde and Satellite Observations Data Set
[9] Signature of Mount Pinatubo eruption aerosols can
clearly be revealed by analyzing the differences between HIRS-2 measurements and radiances simulated by an aero-sol-free radiation model using space-time collocated radio-sonde temperature and humidity measurements as the prime input. This information is directly available from the so-called ‘‘DSD5’’ NOAA/National Environmental Satellite Data and Information Service (NESDIS) collocation archive [Uddstrom and McMillin, 1994a, 1994b]. The space-time window of the collocations is 300 km-3 h. All collocations considered here are cloud free. For more details, see Uddstrom and McMillin [1994a, 1994b] and Che´din et al. [2002]. HIRS-2 brightness temperature measurements are the level 1B original Pathfinder TOVS archive for NOAA 10 (from July 1987 to September 1991) and NOAA 12 (from July 1991 to August 1995).
[10] Because the radiative transfer model used here
(Rapid Radiance Reconstruction Network/3R-N [Che´din et al., 2003]) does not take into account the additional extinction caused by the volcanic aerosols, the difference between 3R-N model-simulated and satellite-observed brightness temperatures should reflect the impact of the presence of the aerosol layer. This additional extinction is actually absorption, since scattering may be neglected at the HIRS-2 wavelengths for Mount Pinatubo aerosols, according to Mie theory results (results not shown; see also Halperin and Murcray [1987]). Following Che´din et al. [2002], differences between simulated and observed brightness temperatures are first centered with respect to their mean, over the whole period (July 1987 to September 1991 for NOAA 10 and July 1991 to August 1995 for NOAA 12), corresponding to their own air mass type (5 classes: tropical, two midlatitude, two polar), as deter-mined from 3R-N. A three-month or twelve-month running mean is then applied, separately for the cases corresponding to nighttime or daytime observations, over land or over sea, and to the three latitude bands: 60 – 20N, 20N to 20S and 20 – 60S. For each channel and each latitude zone, separate statistics of the air mass-centered deviations between modeled and observed bright-ness temperatures for night/land, night/sea, day/land, day/ sea are produced. Hereafter, three-month and twelve-month running centered means are denoted by CM-3 and CM-12, respectively.
2.2. Signature of Pinatubo Aerosols on HIRS-2 Observation Time Series
[11] Four channels, chosen for their different spectral
sensitivities to volcanic eruption aerosols, are considered in the following: HIRS-2 channels 8 at 11.1 mm, 10 at 8.3mm, 18 at 4.0 mm and 5 at 14.0 mm. Channels 8, 10, and 18 are chosen for their high sensitivity to volcanic aerosols
(see Figure 1) whereas channel 5 (14.0mm) is an example of the impact of Pinatubo aerosols on less sensitive channels. Some channels have not been selected because of their instrumental noise or because of their sensitivity to other atmospheric components. For example, channel 9 (9.7mm), located on a strong ozone absorption band, although very sensitive to volcanic aerosols, has not been selected because an accurate simulation of its brightness temperature would require the knowledge of the ozone concentration vertical profile, which is not the case from radiosondes.
[12] For channels 8, 10, 18, and 5, and for the three
zones 20N to 20S, 20 – 60N, 20 – 60S, Figures 2 and 3 show the CM-3 time series for NOAA 10 and NOAA 12 and Figure 4 shows the CM-12 time series for NOAA 12. Expected from the evolution of the Mount Pinatubo aerosol load in the stratosphere, we may see a sharp increase during the three or four months following the eruption (15 – 16 June 1991), then a slow decrease during about three years. Several features on Figures 2, 3, and 4 lead us to assert that the observed signal is due to Mount Pinatubo aerosols: (1) The first is the good agreement between the spatial and time evolution of this signal and in situ or other satellite observations: beginning in the trop-ical zone in June on the three-month running mean case, arrival of the aerosols a few months later in the 20 – 60N and 20 – 60S latitude zones, steady decay of the signal for about two or three years. (2) The second is the strength of the signal, greater than all other noises (statistical noise, satellite or model noise). (The signal-to-noise ratio is about 4; see section 2.3 for more details.) (3) The third is the relative sensitivity of the channels: As expected (see Figure 1), channel 10 shows the greatest deviation in its CM-3 and CM-12, followed by channel 8. Also, channel 5 is not much affected by the aerosols, channel 18 being a bit more sensitive.
2.3. Potential Sources of Error
[13] Figures 2 and 3, in particular, bring into evidence a
seasonal variation of the signal, especially in the Northern and Southern Hemispheres. In the Northern Hemisphere, these seasonal variations have the following behavior: (1) about 0.6 K amplitude; (2) a maximum in summer, from March to August; and (3) a minimum in winter, from November to January.
[14] This approximately seasonal cycle actually results
from the intricate mixing of several influences, as the not accurate enough handling of water vapor in most (if not all) radiative transfer models, the rather modest accuracy of radiosonde measurements of water vapor, and the uneven and variable space and time distribution of the collocations, coupled with the impact of the diurnal cycle (for more details, see Che´din et al. [2002]). This seasonal cycle is removed by going from CM-3 to CM-12 time series, as shown in Figure 4. The NOAA 12 orbital time drift [Christy et al., 2000], which became significant since early 1994 when nearly all the Pinatubo aerosols were gone, is respon-sible for the low positive trend of the time series (more clearly visible in Figure 4).
[15] Because atmospheric profiles over land show a
greater diurnal cycle of temperature and humidity, we only keep over-sea collocations to minimize errors caused by the changes in space and time distribution of the collocations.
Besides, the channels selected (8, 10, 18) may be influenced by surface emissivity, which is more accurately known over sea than over land [Masuda et al., 1988]. Channel 18 being potentially contaminated by solar radiation, only night-time collocations are kept for this channel.
[16] Another potential source of error is the effect of
sulfur dioxide on HIRS-2 radiances, brought into evidence by Prata et al. [2002]. However, the initial SO2cloud was
rapidly converted into aerosols and remained located over a relatively restricted area, in such a way that the time and space averages by the CM-3 and CM-12 series wash out this potential signal.
[17] The noise associated to the Pinatubo aerosol
signa-tures is mostly that of the collocations, through the space-time window, of the radiosonde errors, and of the model. This noise may be reduced by increasing the number of collocations entering the statistics which is larger in the Northern Hemisphere than in the Southern Hemisphere or in the tropics. For channel 10, the error, estimated from the standard deviation of the sample of the differences between model-simulated and satellite-observed brightness temper-atures, varies from about 0.5 K to 0.7 K for the three-month running mean, and from 0.25 K to 0.35 K for the twelve-month running mean. Such values allow for an unambigu-ous detection of the aerosol signal. Obviunambigu-ously, the larger the
number of collocations, the better: We are presently work-ing at increaswork-ing this data set.
3. Modeling the Effect of Volcanic Aerosols on HIRS-2 Radiances
3.1. Radiative Transfer Equation With a Volcanic Aerosol Layer
[18] In the presence of absorbing aerosols, the upgoing
radiance Il at wavelength l emerging at the satellite level
may be written as Il¼ ltml surfð Þt a l surfð ÞBl surfð Þ þ Xn k¼0 Bl½Tð Þk t ml kþ1ð Þtal kþ1ð Þ tm l kð Þt a l kð ÞÞ ð1Þ
where l is the surface emissivity, tlm(k) is the gas
transmittance between the altitude level k and the satellite, tla(k) is the aerosol transmittance between level k and the
satellite, n is the number of considered layers, and Bl[T(k)]
is the black body emission at temperature T and wavelengthl.
[19] In the following, for the sake of simplicity, the
wavelength dependency will be omitted.
Figure 2. Three-month centered running mean deviation (calculated - observed) (CM-3) for NOAA 10: (a) channel 10 (8.3mm), (b) channel 8 (11.1 mm), (c) channel 18 (4.0 mm), and (d) channel 5 (14.0 mm). All collocations are considered here (land, sea, day, night).
[20] The difference I between the radiance emerging at
the satellite level for an free and for an aerosol-loaded atmosphere may be written as
I ¼ BðsurfÞtmðsurfÞ 1 tð aðsurfÞÞ
þX n k¼0 B T½ ð Þk t½ mðkþ1Þ 1 tð aðkþ1ÞÞ ð tmð Þ 1 tk ð að Þk ÞÞ ð2Þ
Provided any other conflicting signal has been properly identified and eliminated, I (or BT in term of brightness temperature) is the theoretical value of CM-3 and CM-12.
[21] The aerosol optical depth, d, can be introduced in
equation (2), through the relationship
tað Þ ¼ exp d kk ð ð ÞÞ ¼ exp X n k sr kð Þdk0 0 ! ð3Þ
where d(k) is the aerosol optical depth between altitude k and the top of the atmosphere,s its absorption cross section, andr(k) its concentration at the altitude k.
3.2. Theoretical Simulation of the Impact of the Aerosols on the Observations
[22] The Pinatubo aerosol model used in this section is
characterized by the monomodal distribution and refractive indices introduced at the beginning of section 2. The aerosol
layer model is made of a triangle-shape layer, located between 14 and 26 km, peaking at 20 km, in agreement with lidar measurements [Vaughan et al., 1994; Chazette et al., 1995], except where other values are specified (e.g., section 3.3.2). [23] The other atmospheric parameters that are needed to
compute brightness temperatures with and without volcanic aerosols, using the radiative transfer equation described above (i.e., ground emissivity and temperature, temperature profile, molecular transmittance profile), are taken from the Thermodynamic Initial Guess Retrieval (TIGR) climatolog-ical data set [Che´din et al., 1985; Chevallier et al., 1998]. TIGR is a climatological library of about 2300 representa-tive atmospheric situations. Each situation is described by its temperature, water vapor and ozone profiles, and its corresponding clear-sky transmittances, radiances and weighting functions for all TOVS sounding channels, com-puted using the fast line-by-line Automatized Atmospheric Absorption Atlas (4A) model in its latest version 2000 (paper in progress).
[24] For more details, the reader may refer to Che´din et
al. [1985] or Scott et al. [1999].
[25] Given a radiosonde measurement collocated with a
satellite observation, the first step consists in searching, among the situations archived in TIGR, the situation the closest to the input situation. The retrieval of the ‘‘closest’’ TIGR atmosphere is based on the minimization of a distance Figure 3. Three-month centered running mean deviation (calculated - observed) (CM-3) for NOAA 12:
(a) channel 10 (8.3mm), (b) channel 8 (11.1 mm), (c) channel 18 (4.0 mm), and (d) channel 5 (14.0 mm). All collocations are considered here (land, sea, day, night).
measuring the proximity of two situations [Che´din et al., 1985; Flobert et al., 1986].
[26] The thermodynamic parameters and precomputed
clear-sky transmittance profiles of the selected TIGR atmo-sphere are then used to compute BT, following equation (2). The sensitivity of BT to atmospheric conditions and aerosol loading and altitude is studied in the next section.
[27] The determination of the closest TIGR atmosphere,
and the calculation of BT are done for all collocated radiosonde profiles. Then, for each latitude zone (20 – 60N, 20S to 20N, 20 – 60S), and for each month, the monthly mean of the simulated BT set is computed to obtain an average value of the theoretical signature of the Pinatubo aerosols.
[28] The last step is the computation of the three-month
and twelve-month running means of the zonally and monthly averaged BT. The brightness temperature differ-ences simulated following this technique are the theoretical values of the CM-3 and CM-12 time series.
3.3. Effects of Atmospheric Conditions, Aerosol Loading, and Altitude on the Simulated Impact of Aerosols on Brightness Temperatures
3.3.1. Effect of the Atmospheric Profile
[29] For a given channel and a given aerosol loading,
the theoretical simulation BT depends on the atmospheric
situation considered. We consider here the 2311 atmo-spheric situations of the TIGR data set. For channel 10, for instance, BT appears strongly dependent on the total water vapor content (Figure 5). The simulations show that BT is larger for humid atmospheres. This depen-dency is the complex result of two physical processes of opposite effects: (1) The contribution of the infrared radiation emitted by the surface (first term of equation (2)) decreases with the molecular surface transmittance, i.e., as the atmospheric water content increases. (2) The contribution of the infrared radiation emitted by the atmospheric layers (second term of equation (2)) is affected by water vapor mainly in the lowest tropospheric layers. For these low altitudes, as there are no volcanic aerosols, the aerosol transmittance is constant. So, this second term is proportional to tm(k + 1) tm(k), the
weighting function, which increases with water vapor content.
[30] The wide range of the results of the simulations
(BT varying from 0 K to 1.6 K) shows that a realistic simulation of the impact of the aerosol on the satellite observations requires knowledge of an accurate description of the underlying atmospheric situation. Considering a fixed standard atmosphere may lead to biased estimates. This explains why our simulation of the BT’s uses situations carefully selected within the TIGR database.
Figure 4. Twelve-month centered running mean deviation (calculated - observed) (CM-12) for NOAA 12: (a) channel 10 (8.3mm), (b) channel 8 (11.1 mm), (c) channel 18 (4.0 mm), and (d) channel 5 (14.0 mm). All collocations are considered here (land, sea, day, night).
3.3.2. Effect of the Aerosol Loading, Optical Depth, and Layer Altitude
[31] Assuming low aerosol optical depth at infrared
wavelengths, d < 0.3, as it is the case for the Pinatubo aerosol layer, equation (2) becomes linear:
I ¼ BðsurfÞtmðsurfÞ Xn k¼0 dð Þk þX n k¼0 B T½ ð Þk tmðkþ1Þ X n j¼kþ1 dð Þj tmð Þk X n j¼k dð Þj ! ð4Þ
which may be simply written, using equation (3), as
I¼Xakd kð Þ ð5Þ
where akis a coefficient depending on the temperature, the
molecular transmittance profile, and the ground emissivity, i.e., not depending on the aerosol characteristics. Figure 6 is a plot of BT as a function of the optical depth for channels 8, 10, and 18. We assume a triangular aerosol layer, spreading from 5 km below the peak to 5 km above the peak, i.e., 10 km deep. For each optical depth, we run 5 simulations with the peak altitude located at 16, 18, 20, 22 and 24 km. The main feature of these plots is the linearity between the optical depth and BT, obviously due to the low aerosol optical depth. Note that BT is not very sensitive to the altitude of the aerosol layer because the clear-sky weighting functions of the channels considered here peak in the low troposphere, and, in consequence, their clear-sky transmittance are roughly constant in the strato-sphere. This is especially true for channel 18 (4 mm), located in a high-transmission atmospheric window. Con-sequently, according to our simulations, a variation of BT must be attributed principally to a decrease in the aerosol optical depth and not significantly to a change in the Figure 5. Variability of the simulated brightness tempera-ture difference, BT, between an aerosol-free and an aerosol-loaded atmosphere at 8.3 mm (channel 10), with total water content, for the 2311 TIGR atmospheric situations.
Figure 6. Variability of the simulated brightness tem-perature difference, BT, between an aerosol-free and an aerosol-loaded atmosphere with the optical depth and the altitude of the peak of the aerosol layer (from 16 to 24 km).
altitude of the aerosols. So, equation (5) may be approximated by
I¼ ad ð6Þ
The coefficient a depends only on the atmospheric and surface conditions and on the channel wavelength, andd is the total aerosol optical depth at the same wavelength. 3.4. Validation of the Aerosol-Loaded Radiative Transfer Simulations Using Lidar Measurements
[32] To compare the CM-3 and CM-12 time series to the
simulation BT, we used the monthly aerosol concentration profiles derived from lidar measurements at the Observa-toire de Haute-Provence (OHP: 44N, 5E) from October 1991 to June 1993 [i.e., Chazette et al., 1995]. Figure 7 shows the aerosol concentration (r(z), in particle number by cubic centimeter) retrieved from lidar measurements at 532 nm. We assume that each profile may represent the aerosol layer above the whole zone 20 – 60N, which is relevant since various measurements show that the layer had become spatially homogeneous over a wide range of lat-itudes in the 20 – 60N band by December 1991 [Jo´nsson et al., 1996; Trepte et al., 1993]. For example, ISAMS retrieved optical depths at 12.1mm show a low variability on the 20 – 60N zone. By the end of 1991, the daily
standard deviation of measurements decreased below 25% of their daily mean (see Figure 8).
[33] Applying the aerosol-loaded radiative transfer
simu-lations described in section 3.2 to the NOAA 12 period, BT is computed for the set of situations obtained by Figure 7. Monthly mean lidar Pinatubo aerosol profiles of the particle number per cm3at OHP, from
Chazette et al. [1995], and their discrete representations on the 4A levels, for four months (November 1991 to February 1992).
Figure 8. Daily mean and standard deviation of ISAMS 12.1 mm optical depth measurements over the 20 – 60N latitude band.
selecting, for each radiosonde-satellite collocation in the 20 – 60N latitude zone, the closest TIGR atmosphere. Then, three-month and twelve-month running means are applied to the simulated BTs. The aerosol signature is simulated for three channels: 8, 10 and 18. As shown in
Figure 9 the agreement is very good with the CM-3 and CM-12 time series. It does not depend significantly on the type of collocations considered (sea only, night and sea only, or all the data), because NOAA 10 and NOAA 12 equatorial local crossing time is 7.30 a.m. and p.m: The Figure 9. Comparison between the experimental (left) CM-3 and (right) CM-12 time series, from
NOAA 12/HIRS-2 measurements and aerosol-free model-simulations using collocated radiosonde profiles, and their theoretical simulations BT (solid line) using TIGR climatological atmospheric profiles and lidar aerosol profiles (three-month and twelve-month running mean, 20 – 60N).
effects of solar contamination and diurnal cycle are not very strong.
[34] For the three-month running mean case, the little
periodic discrepancies are essentially caused by the residual water vapor cycle (see section 2.3). Changes in radiosonde space-time distribution cause other discrepancies. More-over, the water vapor effect is intricately mixed with the aerosol signal: The aerosol mass loading seems to be correlated with the tropopause pressure, which also displays a seasonal cycle [Rosen et al., 1994] and so the deconvo-lution of both effects is very difficult, if not impossible. This effect disappears on the plot of the twelve-month running means (same figure on the right), and the agreement between the experimental and theoretical time series is almost perfect.
4. Infrared Optical Depth Retrieval of Pinatubo Aerosols
4.1. Optical Depth Retrieval Method and Results [35] The optical depth retrieval is based on equation (6).
As shown previously, the effect of altitude, compared with other uncertainties, is negligible, and, in the following, the layer is supposed to spread between 12 and 25 km in the midlatitudes and between 15 and 28 km in the 20N to 20S zone. The inversion algorithm consists in computing BT following the method described in section 3.2, for a reference optical depth (dref = 0.01, value closed to the
expected optical depth), for each channel. The 3- or 12-month running means retrieved optical depth d3 or d12
are given by
d3=12¼
CM3=12 CMcl
BT dref
ð7Þ
where CMclis the value of CM3/12in ‘‘clear’’ (aerosol-free)
conditions. As CM-3/12 are centered mean, we need to evaluate CMclby choosing a particular time when we may
consider the atmosphere as aerosol free to get an aerosol-free level. For NOAA 10, we take the mean of the CM-3 for the period before Mount Pinatubo eruption ranging from July 1987 to June 1990, to avoid seasonal effects, and for NOAA 12, we take the minimum of the CM-3 after the Pinatubo eruption (beginning of 1994). For NOAA 12, because of its orbital time drift, the mean of the CM-3 during the last months is biased, and the aerosol-free level is not very accurately determined.
[36] This retrieval technique is very straightforward and is
applied to three channels (8, 10, and 18), for three latitude zones, for CM-3 and CM-12. The retrieved optical depths are shown in Figure 10 for three infrared wavelengths (4, 11.1 and 8.3 mm). For channels 8 and 10, only sea collocations are used, and for channel 18, only night-sea collocations are used. If the number of collocated measure-ments is lower than 50, then no value is plotted. For the whole period studied here, the maximum aerosol optical depths peak at 0.05 at 8.3 mm, 0.03 at 4 mm, and 0.02 at 11.1mm for the latitude zone 20S to 20N. NOAA 10 and NOAA 12 signals fit themselves quite well in August 1991, except for channel 18 because of the very low number of collocations (contrary to other channels, daily collocations are not used). We should insist on the fact that we obtain
these three infrared optical depth time series with no assumption on aerosol microphysics.
4.2. Comparison With Other Measurements
[37] As mentioned in the introduction, there are very few
measurements of thermal infrared optical depths. Ackerman and Strabala [1994] linked the difference between 11mm and 8.3 mm HIRS-2 channel brightness temperatures with the visible optical depth, and then assuming a refractive index and size distribution, with the infrared optical depth. For the region (0 – 20S and 0 – 30W), for August 1991, they found that the corresponding infrared optical depths for 8.3 and 11.1 mm were 0.03 – 0.06 and 0.014 – 0.02, respectively, depending on the aerosol microphysical and optical model. For the whole latitude band (20S to 20N), we find the following values for the optical depths at the same wavelengths: 0.03 – 0.04 and about 0.013. The agree-ment between the two retrievals is quite good. The differ-ence in the geographical regions considered and their choice of considering only one atmospheric standard profile (tropical from McClatchey et al. [1972]) in their retrieval scheme may explain why our results are slightly lower than theirs. Indeed, the tropical atmosphere of McClatchey et al. [1972] has a total water content of 6.62 cm, which corresponds to a strong aerosol signal on channel 10, according to Figure 5.
[38] To compare our results with previous measurements,
as most observations are made in the visible spectrum, we may compute the particle number concentration. At this level, we need the aerosol size distribution and refractive indices. We use the monomodal distribution and refractive indices given in the introduction of section 2 to get the absorption cross section. Assuming the height and shape of the layer (triangular), we retrieve the particle concentration number, at the layer peak. This should not depend on the wavelength, and the three channels give roughly the same result (not shown). In July and August 1991, we find about 10 to 20 particles cm3at the layer peak (zone 20 – 60N), in agreement with in situ measurements of Deshler et al. [1992, 1993] at Laramie (Wyoming, 41N).
[39] To validate our results with ISAMS optical depths,
we computed the 12.1mm optical depth from channels 8, 10 and 18 retrievals. We assume the same properties for the aerosol as above to compute the aerosol absorption cross-sections at 4, 8.3, 11.1 and 12.1 mm. The ISAMS level 3 aerosol data consist of extinction profiles (in km1) and the corresponding error profiles. The data-to-error ratio distri-bution is bimodal with a minimum around 0.05 (not shown). So, only the measurements with a ratio above this threshold are kept. The comparison between ISAMS and HIRS-2 optical depth for the three latitude zones is quite satisfactory (Figure 11), and differences are probably due to uncertainties in the aerosol size distribution and chemical composition.
[40] Regarding the spatial and time behavior of the
aerosol loading, Figure 10 shows that the transport to the southern latitudes occurred faster than the transport to the northern latitudes, as shown also by the Improved Stratospheric and Mesospheric Sounder (ISAMS) [e.g., Lambert et al., 1993]. In situ measurements also reveal that the optical depth peak occurred in August 1991 for the tropical zone, in November 1991 for the Southern
Hemi-sphere, but not before April/May 1992 in the Northern Hemisphere [Stone et al., 1994]. That is exactly what is observed with our data, and the second peak for the 20 – 60S signal, in July and August 1992 might be a
contam-ination due to polar stratospheric clouds, as also occurred on CLAES data, studied by Mergenthaler et al. [1995] and Lambert et al. [1997]. The twelve-month running mean plots clearly show that Pinatubo aerosols are removed at a Figure 10. Retrieved aerosol optical depths at three infrared wavelengths and for three latitude zones.
(left) Three-month running mean (d3). (right) Twelve-month running mean (d12). From top to bottom:
20N to 20S; 20 – 60N; 20 – 60S. Missing data correspond to months when the number of collocations is lower than 50.
slower rate in the midlatitude region than in the tropics, in agreement with in situ observations of Stone et al. [1994].
5. Conclusions and Perspectives
[41] Covering the time period July 1987 to September
1995, this analysis, at different timescales (seasonal, annual), of the differences between HIRS-2 measurements at 4mm (channel 18), 8.3mm (channel 10), and 11.1 mm (channel 8), and brightness temperatures simulated by an aerosol-free radiative transfer model using collocated radiosonde temperature and water vapor profiles, clearly reveals the ‘‘experimental’’ signature of Mount Pinatubo stratospheric aerosol variations. Not only the strength of the signals, greater than all other noises, and the relative sensitivities of the channels used, in good agreement with what can be expected, but also the space and time evolution of the signatures match quite well in situ or other satellite obser-vations. This experimental signature may also be simulated from a radiative transfer model including the aerosol contribution and used, either in the mode ‘‘with aerosol’’ to simulate the observations, or in the mode ‘‘without aerosol’’ to get the corresponding aerosol-free brightness temperatures. This model, validated against the observa-tions using coincident lidar measurements of the aerosol characteristics, has shown the sensitivity of the HIRS-2 aerosol signal to the atmospheric water vapor content and its relative insensitivity to the altitude of the aerosol layer. As a consequence, and because scattering may be neglected at these wavelengths for this type of aerosol, it may be simply shown that the brightness temperature signal varies linearly with the aerosol optical depth and does not depend on other microphysical properties. On the basis of this linear relationship, the retrieval of the optical depth is straightforward and was applied to the three channels considered here. The retrieved optical depths are found to be in good agreement with the few measurements existing in the infrared. To compare these results with those, numerous, obtained previously in the visible spectrum, we computed the particle number concentration at the layer peak, independently for the three channels considered here. First, this result does not depend on the wavelength, and, for the latitude zone 20 – 60N, the three channels approx-imately give the same result of about 10 to 20 particles by cm3 in July – August 1991, in agreement with in situ observations. Comparison with ISAMS optical depth at 12.1mm is quite satisfactory. The space and time evolution of the aerosol and, in particular, the transport to the northern and southern latitudes also nicely agrees with what has been observed. Moreover, the continuity of the results, obtained from NOAA 10 (ending in September 1991) and NOAA 12 (starting in July 1991), is almost perfect. We may conclude that a dedicated re-analysis of the almost 25 years of HIRS-2 archive offers considerable promise for an improved knowledge of the impact of volcanic aerosols on climate, in particular, through an improved computation of radiative fluxes and budgets. Also, these results strengthen our hope to greatly improve our knowledge of the global distribution of volcanic (among others) aerosols with the new generation infrared vertical sounders like the Advanced Infrared Sounder (AIRS) or the Infrared Atmospheric Sounder Interferometer Figure 11. Comparison between ISAMS aerosol optical
depth at 12.1mm and HIRS-2 aerosol optical depth at 12.1 mm retrieved from channels 8, 10, and 18. Three-month running mean, from top to bottom: 20N to 20S; 20 – 60N; 20 – 60S.
(IASI), both characterized by a much higher spectral resolution and a much larger number of channels.
[42] Acknowledgments. We are grateful to Soumia Serrar for her assistance and expertise in processing HIRS data, to Claudia Stubenrauch and Noe¨lle Scott for their help and comments on the manuscript, and to Nicole Jacquinet-Husson for the Geisa (aerosol) data bank. We thank Didier Tanre´ (Laboratoire d’Optique Atmosphe´rique) and Franc¸ois Dulac (Labo-ratoire des Sciences du Climat et de l’Environnement (IPSL)) for helpful discussions. Thanks are due to Ellsworth G. Dutton (Climate Monitoring and Diagnostics Laboratory, NOAA), who graciously provided us with her Ph.D. report.
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A. Che´din and C. Pierangelo, Laboratoire de Me´te´orologie Dynamique, Institut Pierre Simon Laplace, Ecole Polytechnique, F-91128 Palaiseau Cedex, France. (alain.chedin@lmd.polytechnique.fr; clemence.pierangelo@ lmd.polytechnique.fr)
