Icarus 200 (2009) 118–128

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Icarus

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Mars Global Surveyor Thermal Emission Spectrometer (TES) observationsof variations in atmospheric dust optical depth over cold surfaces

David Horne a,∗, Michael D. Smith b

a University of Missouri, 503 Benton Hall, St. Louis, MO 63121, USAb NASA Goddard Space Flight Center, Greenbelt, MD 20071, USA

a r t i c l e i n f o a b s t r a c t

Article history:Received 4 September 2007Revised 8 September 2008Accepted 3 November 2008Available online 30 November 2008

Keywords:Mars, atmosphereMars, surfaceMars, climateMars, polar caps

The Mars Global Surveyor Thermal Emission Spectrometer (TES) instrument has returned over 200 millionthermal infrared spectra of Mars taken between March 1999 and August 2004. This represents one ofthe most complete records of spatial and temporal changes of the martian atmosphere ever recordedby an orbiting spacecraft. Previous reports of the standard TES retrieval of aerosol optical depth havebeen limited to those observations taken over surfaces with temperatures above 210 K, limiting thespatiotemporal coverage of Polar Regions with TES. Here, we present an extension to the standard TESretrieval that better models the effects of cold surfaces below 200 K. This modification allows aerosoloptical depth to be retrieved from TES spectra over a greater spatiotemporal range than was previouslypossible, specifically in Polar Regions. This new algorithm is applied to the Polar Regions to show theseasonal variability in dust and ice optical depth for the complete temporal range of the TES database(Mars Year 24, Ls = 104◦, 1 March 1999 to Mars Year 24, Ls = 82◦, 31 August 2004).

© 2008 Elsevier Inc. All rights reserved.

1. Introduction

Mars Global Surveyor (MGS) has conducted systematic mappingof Mars from 1 March 1999 (Mars Year 24, Ls = 104◦) to 31 August2004 (Mars Year 27, Ls = 81◦). Where Mars Year (MY) is definedas in Clancy et al. (2000), with the beginning of MY1 at 11 April1955. The Thermal Emission Spectrometer (TES) mounted on thenadir deck of the spacecraft, has amassed a database of more than200 million detailed thermal infrared spectra. These spectra can beemployed to determine both atmospheric and surface properties(e.g., Conrath et al., 2000; Christensen et al., 2001; Bandfield et al.,2000; Smith, 2004).

The long observational history and systematic global coverageof the TES dataset presents an invaluable resource for continuouslong duration surveys of martian atmospheric variability. A greatdeal of work has been conducted in this field and reliable methodshave been previously developed to retrieve key atmospheric quan-tities such as atmospheric temperature (Conrath et al., 2000), dust(Smith et al., 2000a; Smith, 2004), ice aerosol optical depth (Pearlet al., 2001), and water vapor abundance (Smith et al., 2002).

Using these algorithms, TES has been utilized to study theseasonal and spatial atmospheric variability in the martian at-mosphere. Despite the near-continuous monitoring of the martianatmosphere, prior studies of the seasonal and spatial variation of

* Corresponding author.E-mail addresses: horneda@umsl.edu, dhorne@umsl.edu (D. Horne).

0019-1035/$ – see front matter © 2008 Elsevier Inc. All rights reserved.doi:10.1016/j.icarus.2008.11.007

aerosol optical depth have been limited to regions and times hav-ing a large and positive thermal contrast between surface andatmosphere, which limited coverage to spectra obtained over sur-faces with temperatures above about 210 K. Consequently, therecord of aerosol retrievals is incomplete at higher latitudes. Theseasonal cycle of advancing and retreating polar caps has effec-tively prevented the retrieval of a continuous temporal dataset foraerosol optical depth outside a mid latitude band of around ±30◦ .These limitations are effectively illustrated in Smith (2004).

Visual wavelength imaging and the available infrared observa-tions of Mars along with circulation models (Pollack et al., 1990)have shown that dynamic weather activity takes place over bothpoles during winter (Cantor et al., 2001). The thermal conditionsof the atmosphere in the vicinity of the poles are influenced bythe autumn/winter advancement and subsequent spring/summerregression of the polar cap. Both the MGS MOC instrument andHST visual cameras have observed dust activity coincident with theexpansion and regression of the polar cap (James et al., 2005). Thisis one aspect of the seasonal dependence of dust and ice aerosoloptical depth (Pankine and Ingersoll, 2004).

Here, we present modifications to the standard TES aerosolretrieval algorithm that are useful for observations over cold sur-faces. An overview is included of the seasonal and spatial variationof dust and water ice optical depth activity over the cold (winterseason) martian poles for the complete TES database covering theperiod 1 March 1999 (Ls = 104◦ , Mars Year 24) to 31 August 2004(Ls = 81◦ , Mars Year 27). This presents a nearly complete record ofpolar aerosols for the almost three Mars years of TES observations.

TES: Atmospheric dust optical depth measurement over cold surfaces 119

2. TES data and the retrieval algorithm

Only nadir pointing observations have been utilized in thiswork. Nadir observations not only represent the bulk of the TESdatabase but also form a systematic map of the planet as thespacecraft progresses in its orbit. During its mapping phase MGSprovides two complete planetary maps per day, one at approxi-mately 1400 h local time and another at 0200 h. Retrievals areperformed for each 3 × 2 pixel TES footprint giving a spatial reso-lution of either 9 × 10–20 km (Smith et al., 2000a).

The algorithm presented here is based on the retrieval of Smithet al. (2000a) and is designed primarily to retrieve atmosphericdust optical depth. Modifications have been implemented in thisalgorithm to obtain reliable optical depths over areas and timeswhere surface temperature is below 210 K.

2.1. Data processing procedure

In this work we restrict our retrievals to daytime (∼1400 h lo-cal time), nadir-geometry spectra where the surface temperatureis less than 210 K. This effectively limits the spatial coverage tomid-latitudes and higher in the winter hemisphere. The opacityretrieval algorithm is composed of a radiative transfer algorithmand a least-squares fitting routine. The column integrated aerosoloptical depth of dust and ice is determined by values providingthe best least squares fit between the computed and the observedradiance spectrum.

2.2. Radiative transfer

The observed radiance of Mars as observed by TES, may be de-scribed in terms of an absorbing atmosphere approximation. Thiscan be written in a simplified form as

I(υ) = ε(υ)B(Tatm,υ)e−τ dτ +

τ0∫

0

B(Tsurf,υ)e−τ0(υ), (1)

where B(T ,υ) is the Planck function for a given temperature, T ,and frequency, υ,ε(υ) is the surface emissivity as a function offrequency, and Tsurf and Tatm are surface and atmospheric tem-peratures respectively. Following Smith et al. (2000a, 2000b), weneglect scattering, assume a well-mixed aerosol and set surfaceemissivity to unity. The integration is performed from the space-craft to the surface of the planet for a specific value of column-integrated aerosol optical depth τ0(υ). The total aerosol opticaldepth is expressed as a linear combination of contributions fromdust and ice. Using given spectral shapes of dust ( fdust), ice ( f ice),their respective scaling factors (Adust) and (Aice), the total opticaldepth can be defined by

τ0(υ) = Adust fdust(υ) + Aice f ice(υ). (2)

2.3. Inputs and assumptions

In order to compute the expected radiance using Eq. (1), at-mospheric temperature and surface temperature are required asinputs to the radiative transfer code. We use atmospheric tem-peratures that were previously obtained by the TES team usingan inversion of the 15-micron CO2 band along with correspondingsurface temperatures found through the observation of equivalentbrightness temperature at 1300 cm−1 (Conrath et al., 2000).

The assumption that aerosols are well-mixed with the CO2 gasof the atmosphere, allows atmospheric pressure to be directly re-lated to total aerosol optical depth. TES dust opacity data scale wellwith surface pressure and emission angle (Smith et al., 2000b);hence dust optical depth is considered proportional to the amount

of gas along the line of sight. Dust storm vertical structure ob-servations (Jaquin et al., 1986; Smith et al., 1997) support thisassumption. The optical depth at a given altitude is therefore con-sidered proportional to the atmospheric pressure at that point.This proportionality between optical depth and pressure allows theintegration over optical depth to be replaced with an equivalent in-tegration over atmospheric pressure, simplifying the evaluation ofEq. (1).

The well-mixed aerosol assumption does not hold as well forwater ice clouds and may result in inflated ice opacity values.Numerical experiments by Smith (2004) show absorption opticaldepths for TES nadir-geometry observations typically underesti-mate extinction optical depth by about 30% for dust and about50% for ice. Total error budgets differ from those stated in pre-vious works due to the nature of the modifications and these areaddressed in Section 2.5.

We use the standard spectral functions ( fdust, f ice) for dust andice (Smith, 2004) and assume that these functions accurately de-pict the spectral characteristics of dust and ice particulates. Thesespectral functions are assumed to be independent of time or posi-tion.

It is assumed that the majority of brightness temperature spec-tra retrieved over cold surfaces will exhibit a distinctive positiveslope to varying degrees and that this gradient can be accuratelymodeled by the fitting of three factors; initial surface temperature(T 1), a second temperature (T 2) equal to T 1 + 10 and a mixingfactor ( f ) with a value between 0 and 1.

2.4. Modification for cold surfaces

Radiance spectra in this study are computed using TES surfacetemperature estimates as a starting point in the radiative trans-fer code. Spectra retrieved over low temperature surfaces oftenexhibit an equivalent brightness temperature that increases withfrequency. Fig. 1 demonstrates the change in the relative gradi-ent of the brightness temperature spectra with diminishing surfacetemperature. The effects of changing surface–atmosphere tempera-ture contrast are also evident in this figure. Temperature retrievalstaken from areas of Tsurf = 110 to 180 K, are shown along withtheir associated radiance spectra.

The spectral gradient is believed to be the result of a rangeof surface temperatures in the field of view of the detector. Forexample, a single TES pixel may observe a mixture of both frost-covered and (much warmer) bare ground. We model this effect byexpanding the first term in Eq. (1), giving Eq. (3) which describesthe radiation from the surface in terms of a mixture of two surfacetemperatures.

B(ν,Tsurf) = f B(ν,T 1) + (1 − f )B(ν,T 2). (3)

The values chosen for each surface temperature (T 1 and T 2)and those of the mixing fraction ( f ), accurately model the varietyof spectral gradients observed in cold surface retrievals. A numberof combinations of mixing factor and temperature differential canresult in a similar continuum slope. It is therefore difficult to ascer-tain which particular variable (temperature differential or mixingfactor) or value thereof, is responsible for inducing the observedgradient. This prevents accurate conclusions from being drawn asto the true physical nature of the surface. The introduction of astarting temperature indicative of polar cap ice may produce amore physically meaningful solution and this is being consideredfor any follow up observations. Simulations were conducted to testthe reliability of the fitting method; in each case a different vari-able was altered until a best fit to the data was achieved for thegreatest number of cases. In the first case a variety of values forT 1 were utilized while holding T 2 and f constant. In the second

120 D. Horne, M.D. Smith / Icarus 200 (2009) 118–128

Fig. 1. Typical surface radiances with corresponding brightness temperature spectra over surfaces from 110 to 220 K. Note the change in the relative slope of the brightnesstemperature spectrum. The varying depth of the 15 μm atmospheric CO2 feature compared to the continuum level is also an indicator of the surface–atmosphere thermalcontrast.

Fig. 2. The effect of varying the distribution of the radiance spectrum on the corresponding brightness temperature. The mixing ratio ( f ) is applied to simulated radiancemeasurements and varied from 0.02 to 0.1 (panel 1). Corresponding brightness temperatures resulting from the application of Eq. (3) using T 1 = 140 K and T 2 = T 1 + 10 areseen in panel 2.

case T 2 was altered while T 1 and f remained invariant. Finally,f was varied while T 1 and T 2 were held constant.

By varying f between 0 and 1, the radiance distribution canbe successfully altered to model the observed temperature gra-dient. Fig. 2 shows the effect of applying values of f from 0.02to 0.1 to the radiance spectrum; note the relative change in thegradient of the brightness temperature spectrum as f increases.Simulations such as these demonstrate the sensitivity of the radi-

ance spectrum to the application of the weighting factor and thecorresponding change in the slope of the brightness temperaturespectrum.

Experiments were conducted to match the form of real bright-ness temperature spectra, in order to determine a suitable math-ematical expression for the form of the continuum slope. Fig. 3shows a sample of the intermediate steps cycled through by thealgorithm during tests to determine the value of the mixing frac-

TES: Atmospheric dust optical depth measurement over cold surfaces 121

Fig. 3. Sample results showing several iterations run through by the modified algorithm to match the characteristic slope of a single, real brightness temperature spectrumobtained over a cold surface. Radiance spectra corresponding to each brightness temperature spectrum are presented in panel (a).

tion which best fits the observed spectrum for two temperatures10 K apart. Models such as these were utilized to examine theform of the brightness temperature spectrum and its response tochanges in radiance, temperature and the application of f . Sim-ulations indicated that a better fit was achieved and convergencetimes were reduced by varying only the value of f . Several hun-dred cases were visually inspected as per Fig. 3 for a range ofdifferent starting temperatures from 110 to 220 K, in order toensure the accuracy of the fitting and minimization routine. Theseverity of the spectral gradient is determined primarily by thevalue of the mixing factor f . Values of f were generally between0.85 and 0.95 for surface temperatures from 140 to 180 K, as tem-peratures drop so the value of f tends more toward unity. Aftercompensation for the spectral slope, the determination of dust andice optical depth values was carried out as per the Smith (2004)algorithm.

Powell’s method of minimization for multiple dimensions wasutilized to fit five separate parameters: two surface temperatures(T 1, T 2), the mixing ratio, f , and the reference scaling factors fordust and ice optical depth (Adust and Aice). Initial testing involvedthe fitting of each parameter individually; the fits for each param-eter were then compared to the final result for all five variablesto check for inconsistencies introduced by the fitting of multiplefactors.

2.5. Uncertainties in the retrieved values

The retrieval algorithm presented here is based on that ofSmith et al. (2000a) and is subject to many of the same sourcesof error with additional considerations relating to the new mod-ifications. Instrumental noise is stated to be on the order of2.5 × 10−8 W cm−2 sr−1 cm−1 with instrumental calibration effectscontributing around 1×10−7 W cm−2 sr−1 cm−1 (Christensen et al.,2001). The effect of varying solar angle is minimized by exclusivelyutilizing afternoon nadir pointing spectra. The above effects con-tribute to a cumulative error of around 0.02 in the final opticaldepths. Errors introduced through the application of fixed spec-tral shapes remain at around 0.01–0.02. Uncertainty in the startingtemperature (T 1) is similar to that of the original method for tem-peratures over 140 K and lies at around 3 K.

Uncertainty in initial surface temperature (T 1) is potentiallythe greatest source of error. Inaccuracies in the determination ofT 1 and the subsequent value of T 2 may result in over or underestimation of the spectral gradient and the loss of physically realis-tic retrievals. Surface temperature determinations tend to becomemore unreliable over colder surfaces due to the presence of in-creased noise. In order to determine the accuracy of these values,surface temperatures retrieved from the algorithm were comparedto existing temperature values from the TES database as discussedbelow.

122 D. Horne, M.D. Smith / Icarus 200 (2009) 118–128

Fig. 4. A comparison of key values retrieved by the modified algorithm to simulated and TES retrieved data for 2 × 103 observations with surface temperatures ranging from110 to 180 K. (a) Correlation between retrieved surface temperatures and estimates from the TES database. (b) Q compared to surface–atmosphere temperature contrast.(c) Distribution of Q relative to retrieved optical depth.

Experimentation on multiple spectra showed that a value forT 2 of T 1 + 10 was sufficient to model the general slope of thebrightness temperature spectrum (and the corresponding radiancedistribution). Using this value of T 2 as a starting point, the weight-ing factor ( f ) is applied between T 1 and T 2 to match the spectralslope. This method of selecting T 2 can result in under and overestimates of the initial slope by as much as 20% in extreme caseswhere the observed slope is very steep or very shallow. This usu-ally occurs over very cold surfaces, or in cases where thermalcontrast is very low and these results are usually discarded. De-spite the initial uncertainty that may result from poor temperaturedetermination, an accurate estimation of the spectral slope can stillbe determined, as f can be adjusted until a match to the observedspectral gradient is achieved (see Fig. 3).

Fig. 4a shows a comparison of around 20 × 103 temperaturevalues retrieved from the TES database with those determinedby the modified retrieval algorithm for the same time and loca-tion. These temperatures are in agreement to within 10% for the180–220 K temperature range, increasing drastically to around 35%for temperatures less than 130 K. Radiance spectra below 120 K

tend to become noise dominated, preventing accurate tempera-ture determinations. As such, optical depth retrievals over surfaceswith temperatures of 120 K or less are considered to be unreli-able.

The reliability of the algorithm was continually tested; syntheticradiance spectra were generated for a range of optical depths andcompared to observational results. The difference between com-puted and observed spectra was used to produce a factor (Q ), seenin Eq. (4) which demonstrates the quality of the fit to real data.

Q = (Icomp − Iobs)2. (4)

Fig. 4b shows the variation in Q with the magnitude of thesurface–atmosphere temperature contrast for the 20 × 103 pointsseen in panel (a). Fig. 4c demonstrates the distribution of Q val-ues with respect to retrieved dust optical depths. The relative errorbudget for each temperature range along with the correspondinguncertainty in dust optical depth is summarized in Table 1.

The higher relative error budget for dust optical depth com-pared to surface temperature may be due to the increased noiseinherent to spectra over low temperature surfaces. Inaccurate sur-face temperature determination may result in poor estimation of

TES: Atmospheric dust optical depth measurement over cold surfaces 123

Fig. 5. Zonally averaged seasonal and latitudinal variations in polar surface temperature, ice and dust aerosols for the period MY24–MY27. Time is shown in terms of MarsYear (MY) (top line), the end of each year defined by a dashed vertical line. The horizontal axis shows cycling Ls values for each year. The vertical axis shows latitude from

◦ ◦ ◦ ◦

−90 to 90 . Data are binned into 2 × 2 intervals.

Table 1Error budget for four temperature bands from Tsurf = 120 to 220 K with corre-sponding uncertainties in derived dust optical depth.

Temperature (K) Relative error: Tsurf Relative error: Dust optical depth

180–220 0.10 0.12150–180 0.15 0.18135–150 0.20 0.25120–135 0.28 0.34

temperatures in the 90 successive atmospheric pressure levels runthrough by the algorithm. As the optical depth retrieval is depen-dent on pressure, cumulative errors in atmospheric temperatureserve to increase the uncertainty in optical depth.

Neglecting uncertainties in T 1 and T 2, error in the value off is limited by the accuracy of the fitting and minimization rou-tine used to determine the gradient and should remain relativelyconstant with an upper limit of 0.01. As a control, numerical sim-ulations were conducted using fabricated inputs to ascertain theoverall effect of modifying the radiance distribution. The resultsof the algorithm for simulated data were compared to actual re-trievals in order to ascertain the overall reliability of the gradientfitting method. Figs. 2 and 3 are examples of these types of simu-lations. While the mixing factor is capable of accurately matchingthe spectral gradient, extreme under or over estimates in startingtemperature may serve to artificially inflate or deflate tempera-ture estimates. This effect tends to occur in noisy spectra retrievedfrom very low temperature surfaces (Tsurf = 120–135 K). Compar-ison of observations to simulated spectra minimizes this effect. Ifthe difference in the temperature derived from the simulated spec-tra compared to the observed is significant (<10 K), then it is likelywe are dealing with an anomalous result. The judicious use of Qfactors allows for the effective screening of poor retrievals over ar-eas with very cold surfaces or very low temperature contrast.

Fig. 4b demonstrates the reliability of the algorithm (Q ) com-pared with atmosphere–surface temperature contrast (�T ). Thealgorithm provides reliable retrievals at high temperature contrastswith simulated and observed spectra exhibiting a minimum dif-ference at around �T = 20–50 K. Differences in the two spectraincrease steadily with diminishing temperature contrast, reachingQ = 10 by �T ≈ 8. While low temperature surfaces can be suc-cessfully accounted for, the algorithm cannot compensate for theeffects of low temperature contrast. The reliability of the opti-cal depth retrieval is demonstrated in Fig. 4c. Reliable dust opti-

cal depths span a narrow, physically realistic range of Q valuesfrom 1 to 10. The optical depth distribution steadily increasesfrom around Q = 6–10, drastically increasing from Q = 10–20. ByQ = 14 optical depth estimates span a wide range (2.0 and above)indicating a high level of uncertainty in any one optical depth re-trieval. While retrievals with Q factors greater than 5 may be con-sidered on a case by case basis, results with a Q value in excess of10 are considered to be physically unreliable. Optical depths withhigher Q values are generally associated with areas/times possess-ing very cold surfaces (T < 120 K) or very low thermal contrast,and will likely exhibit a very poor signal to noise ratio.

This algorithm is designed to search for increases in dust ac-tivity and so the assumption that aerosols are well mixed maynot hold as well for the case of water ice clouds which form atdiscrete altitudes where temperatures are favorable for condensa-tion (Pearl et al., 2001). This effect may result in overestimationof water ice optical depth by around 20% (Smith, 2004). Ice opti-cal depth results will also be subject to the same sources of erroras the dust retrieval. Specific uncertainties relating to ice retrievalsare difficult to judge without increased knowledge of a specific iceclouds’ physical characteristics, which is beyond the scope of thisproject. Despite these limitations, ice opacities have been includedfor completeness and comparison to dust optical depth results (seeFig. 5). Relative increases in ice optical depth can be observedwithin the context of the results presented here.

The assumption of unit surface emissivity remains a sourceof uncertainty; lack of compensation for this effect may increasenoise levels for cold surface retrievals. This may result in overesti-mates of optical depth for areas with lower emissivity by as muchas 10% in the most extreme cases. It is the intention of this studyto evaluate the overall contribution of the spectral gradient to thefinal retrieved optical depth, neglecting this correction allows theoverall effect of the gradient modification on optical depth valuesto be more clearly understood. Downwelling effects may also con-tribute to the error budget; this effect is strongest at the height ofwinter when high albedo ice/snow coverage is at its peak. It hasbeen shown that airborne dust acts to reduce the downwelling so-lar flux (Savijärvi et al., 2005). Consequently, for instances wheredust optical depth is appreciably high, the effects of downwellingradiation are minimized, any error contributed by this effect islikely to be small.

124 D. Horne, M.D. Smith / Icarus 200 (2009) 118–128

Fig. 6. Zonally averaged seasonal and latitudinal variations in surface temperature, ice and dust aerosols for all latitudes and all times for the period MY24–MY27. Time isshown in terms of Mars Year (MY) and Ls for each year. Each Mars Year is defined by a dashed vertical line. Latitude is shown from −90◦ to 90◦ . Data are binned into2◦ × 2◦ intervals. The color scale has been adapted to match that of Smith (2004) in order to demonstrate continuity of polar data with lower latitudes.

3. Results

The modified algorithm was used to retrieve dust and ice op-tical depth values for each daytime spectrum in the TES databasewith a surface temperature below 210 K. Optical depth for dustis given in terms of a reference frequency of 1075 cm−1, whilewater ice optical depth is given at 825 cm−1. Fig. 5 shows thezonally averaged seasonal and latitudinal variations in surface tem-perature over the North and South poles along with correspondingdistributions of dust and water ice optical depth. The data shownhave been subject to spatial averaging in order to minimize noiseeffects and any spurious values. Data were subsequently binned2◦ × 2◦ in latitude and Ls. Repeating seasonal patterns in dust andice optical depth are evident in both hemispheres over all sur-veyed Mars years. Due to insufficient thermal contrast betweenthe surface and atmosphere, areas and times exist over whichthe algorithm is incapable of retrieving optical depth. This oc-curs along the transition region between low-latitude areas whereTatm < Tsurf and higher latitude polar regions where Tatm > Tsurf.This results in the areas of missing data seen in Fig. 6, whichshows the results of Smith (2004) combined with those of themodified algorithm. This provides an almost complete record ofdust optical depth from Ls = 90◦ , MY24 to Ls = 90◦ , MY27. Exceptfor the spatiotemporal discontinuities resulting from poor tem-perature contrast, the agreement between the new polar opticaldepth measurements and those of the Smith (2004) survey canbe clearly seen. This continuity of atmospheric features betweenthe 2004 dataset and new results demonstrates the reliability ofthe dust optical depth retrievals. This continuity is most appar-ent during the periods of high dust activity from Ls = 180◦ to 270◦each year, and is especially evident during the 2001a dust storm inMY25.

3.1. North polar observations

3.1.1. Mars Year 24Dust optical depth increased from 0.1 to 0.4 at around Ls =

180◦ . This activity initially stretched from 40◦ to 70◦ N latitude un-til Ls = 210◦ , when dust activity formed two discrete bands from43◦ to 47◦ and 60◦ to 70◦ N latitude. An annulus of low dust op-

tical depth values (0–0.1) remained in a latitudinal zone from 47◦to 60◦ N latitude until a resurgence in dust activity occurred fromLs = 295◦ to 344◦ . These annuli, while evident in Figs. 5 and 6are most clearly seen in line 1 of Fig. 7a at Ls = 249◦ and 292◦ .At around Ls = 310◦ dust optical depth increased between 65◦ to90◦ N latitude, reaching optical depths in excess of 0.5. Dust op-tical depths rapidly subsided to a low (∼0.1) level after Ls = 10◦in MY25. Surface temperatures were relatively uniform over theNorth pole during MY24, with a clear latitudinal cutoff in temper-ature. Higher latitudes (65◦–90◦ N) exhibited temperatures downto 120 K, with slight warming from around Ls = 310◦–356◦ . Whiletemperatures immediately south lay at around 140–150 K. This sta-ble environment may have contributed to the general absence ofNorth polar dust storms in MY24 compared to successive years.

3.1.2. Mars Year 25Mars Year 25 exhibited similar behavior to MY24, however dust

optical depth values began to increase earlier at around Ls = 160◦ .Dust optical depths over 70◦ N were around 60% higher for theperiod, Ls = 180◦–310◦ in MY25 compared to MY24. This in-crease is most likely due to dust raised by planet encircling duststorm 2001a which took place from Ls = 185◦ to 215◦ (southernspring) in MY25 (Smith, 2004). High dust optical depths were morewidespread around the vicinity of the pole and extended south tolower latitudes.

Dust storm 2001a originated at low southern latitudes in theHellas Planitia region. The focus of this study was to observe dustoptical depths over cold surfaces, which appear predominantlyover the autumn–winter periods. Consequently, no observationswere made southward of around 40◦ N during peak dust stormactivity, as these areas possessed temperatures outside the rangeof this investigation.

As in MY24, the region of high dust optical depth separatedinto two latitudinal bands, creating an annulus of low dust op-tical depth (0–0.1) which lasted from Ls = 220◦ to 310◦ . Dustactivity increased at Ls = 310◦ , by Ls = 345◦ dust optical depthreached 0.5 and above with highest values near the pole. Particu-larly strong activity was observed between 60◦ and 90◦ N latitudefrom Ls = 345◦ to 10◦ (MY26). Dust optical depth had mostly re-turned to the 0–0.1 level by Ls = 35◦ (MY26). Surface temperatures

TES: Atmospheric dust optical depth measurement over cold surfaces 125

(a)

(b)

Fig. 7. The spatial distribution and values of dust optical depth (blue = 0, red � 0.5) for regular intervals throughout the martian seasonal cycle in the Northern hemisphere(a) and Southern hemisphere (b) from MY24–MY27. (a) Note the seasonal repeatability present in each year. MY24 and MY25 both show pronounced annular bands oflow dust optical depth from Ls = 249◦ to 292◦ . (b) Note the seasonal repeatability present in each year and the absence of the banding effects observed in the Northernhemisphere.

126 D. Horne, M.D. Smith / Icarus 200 (2009) 118–128

in MY25 were not as uniform as MY24 which possessed a clear lat-itudinal delineation in temperature. Increased temperatures weredetected at high latitudes, with pockets of isolated warming occur-ring over the upper pole at Ls = 261◦ and 323◦ . This warming wascoincident with strong dust detections.

3.1.3. Mars Year 26–27Higher dust optical depth was observed during MY26 between

55◦ and 90◦ N latitude than during previous years in the same sea-son. Dust optical depth values of 0.5 to 0.8 were observed betweenLs = 140◦ and 190◦ at 68◦ to 90◦ N latitude. Dust optical depth in-creased at Ls = 190◦ reaching values of 0.5 and above over a widelatitude range of 40◦ to 90◦ N with the highest optical depth val-ues poleward of 60◦ N latitude. As in previous years, a band of lowdust optical depth formed between 45◦ and 60◦ N latitude, ceasingto exist at Ls = 305◦ when dust levels increased. This band of lowdust optical depth differed slightly from previous years in that thefeature was more constrained in its latitude range and was shorterlived. Dust optical depth at 60◦–90◦ N latitude was much higherand longer lived than previous years, with values in excess of 0.5observed almost consistently from Ls = 193◦ to 29◦ (MY27). Dis-tinct increases in dust optical depth took place at 60◦ to 90◦ Nlatitude between Ls = 266◦ and 313◦ , and between Ls = 0◦ and29◦ (MY27). Dust optical depth fell significantly to around 0.1 byLs = 35◦ (MY27). Surface temperatures in MY26 followed a simi-lar trend to the preceding year, with isolated warming occurringat around Ls = 271◦ and 347◦ . Again, these areas experienced in-creased dust activity.

3.2. South polar observations

3.2.1. Mars Year 24–25The TES dataset begins at Ls = 104◦ for MY24 corresponding to

mid-winter in the southern hemisphere. For the period Ls = 104◦to 231◦ dust optical depth remained relatively low, with a slightincrease to a value of 0.2 observed as surface temperatures beginto climb between Ls = 200◦ and 230◦ .

Shortly after Ls = 0◦ , MY25, dust optical depth increasedrapidly, reaching values of 0.5 and above, extending northwardreaching from the pole to 45◦ S latitude by Ls = 95◦ . Dust opticaldepth was generally high poleward of 55◦ S at Ls = 138◦ and 216◦ .Dust activity increased near the pole at the conclusion of the yearbetween Ls = 345◦ and 360◦ . This study concentrated on cold sur-faces which are minimal during the summer season, meaning theeffects of the MY25 dust storm were largely excluded. Surface tem-peratures were generally well defined with temperatures of 120 Kreaching around 54◦ S at the height of winter. Some warming wasevident from Ls = 27◦–98◦ (MY25) coincident with increased dustactivity.

3.2.2. Mars Year 26–27Once again, dust optical depth began to increase shortly af-

ter Ls = 0◦ , MY26 reaching values of 0.5 and above, poleward of50◦ S latitude between Ls = 340◦ and 105◦ . Aside from a slight de-crease in dust activity between Ls = 0◦ and 90◦ , dust optical depthremained near 0.5 poleward of 50◦ S latitude, later declining be-tween Ls = 175◦ and 190◦ . During the period Ls = 25◦ to 100◦ , thedust activity observed in MY26 was generally weaker and shorterlived than MY25.

Dust optical depth increased rapidly at Ls = 0◦ (MY27) reachingvery high values and expanding northward to a latitude of 65◦ S.Dust optical depth fell at around Ls = 30◦ (MY27) and while ele-vated poleward of 50◦ S latitude dust opacity remained lower thanprevious years. As in MY25, surface temperatures climbed over thepole at the beginning of the year, however this occurred slightly

earlier in MY26 at around Ls = 10◦ , this increase in temperaturewas spatially larger and longer lived than the preceding year.

Some smaller scale warming was also evident later in the yearat around Ls = 175◦ . Comparison to the previous year is difficultdue to missing data in MY25. MY27 temperatures were only avail-able from Ls = 0◦–90◦; however polar temperatures appeared tobe colder and lacked the sporadic warming evident over this pe-riod in previous years.

4. Interannual comparison

4.1. Northern hemisphere

Dust optical depth increased over the north pole for the latterpart of each martian year and early into the subsequent year fromLs = 320◦ to 20◦ , with high optical depths recorded from around60◦ to 90◦ N over this time. North polar dust detections under-went a further increase each year at around Ls = 200◦ . High dustoptical depths extended from the pole to a maximum elongationof about 40◦ N in all three years. Dust optical depths dropped tovery low levels in an annulus between 45◦ and 60◦ N at aroundLs = 230◦ each year (slightly later in MY25), while higher dustopacities were maintained to the north and south of this region.Fig. 7a shows the spatial variation of dust optical depth retrievalsin the northern hemisphere for specific Ls ranges over three com-plete Mars years.

The North pole showed lower overall dust activity in MY24 thanin subsequent years. Both MY25 and 26 exhibited very strong dustresponses poleward of 60◦ N latitude between Ls = 160◦ and 240◦ ,with dust optical depth remaining high until Ls = 360◦ . DuringMY26, discrete increases in dust optical depth were more promi-nent in both strength and duration, with two such increases takingplace at Ls = 197◦ and 266◦ . High dust optical depths also ex-tended further south to 40◦ N during MY25 and 26 compared toMY24. A common feature across all surveyed years was an increasein dust opacity around the onset of northern spring (Ls = 0◦) pole-ward of 50◦ N. Dust optical depths for this Ls in MY26 werearound 10% higher and longer-lived than preceding years. As thepolar cap is receding at this time, thermal gradients may form be-tween cold polar ice and the comparatively warmer bare ground.This may result in a front that lifts dust across the seasonal po-lar cap (James et al., 1999). This is supported by the coincidence ofelevated dust optical depth with a significant variation in temper-ature; this correlation is most pronounced in Fig. 5.

While North polar dust optical depths were elevated in MY25due to the increased activity of the 2001a dust storm, warmingover this period was not as widespread as Viking atmospherictemperature measurements (Jakosky and Martin, 1987). Localizedincreases in surface and atmospheric temperature were detectedin areas exhibiting particularly high dust optical depth. Here weconcentrate on the effects of surface temperature, further analysisand mapping of atmospheric temperatures may be the focus of fu-ture work. The greatest increases in dust optical depth around theNorth polar cap occur along the Tsurf ≈ 140–150 K (dark blue-lightblue) boundary from 60◦–70◦ N, Ls ≈ 190◦–300◦ and 40◦–50◦ N,Ls ≈ 193◦–350◦ at the Tsurf ≈ 150–160 K (light blue-green) inter-face. These areas correspond to the upper and lower latitudinallimits of the clear annular bands. The increased dust storm activityalong the cap edges at Ls ≈ 193◦–350◦ is confirmed by prior MOCobservations (Cantor et al., 2002). Similar correlations with tem-perature contrast are also seen around the south pole along theinterface of the 140–150 K (dark blue-light blue) regions at around45◦–85◦ S, Ls = 0◦–90◦ .

Possible detections of water ice over the North Polar Regionwere widespread between 35◦ and 80◦ N latitude with opticaldepths typically 0.3–0.5. Water ice optical depth appeared to drop

TES: Atmospheric dust optical depth measurement over cold surfaces 127

to negligible levels in each year between Ls = 10◦ and 70◦ . Higherrelative ice optical depths were observed at 37◦ N each year fromLs = 300◦ to 330◦ . Bands of low ice optical depth also formed from45◦ N and 60◦ N from Ls = 220◦ to 310◦ , however these bandsappeared to encompass a slightly greater latitudinal and temporalrange than corresponding bands in dust optical depth.

4.2. Southern hemisphere

Distinct seasonal cycles in dust activity were also observed inthe southern hemisphere; however the pronounced annular fea-tures seen in the North were not evident in the South. South polardust activity also lacked the shorter spatial and temporal variationsseen in the North, instead undergoing distinct increases and de-creases coincident with the seasonal cycle. In each surveyed year,early summer (Ls = 300◦–350◦) dust optical depths remained be-low 0.1, rapidly increasing at around Ls = 350◦ each year. Thisregion of high dust optical depth steadily expanded northward,reaching around 50◦ S latitude by Ls = 90◦ in the subsequent year.After Ls = 90◦ , dust optical depth values fell and began to recedesouthwards, dust activity reduced in size and became more spo-radic. Dust optical depth values finally subsided to a backgroundlevel (>0.1) at Ls = 180◦ in each year. MY25 showed higher overalldust opacities in the south compared to other surveyed years. Dustopacity near the South pole was significantly higher for MY25 thanMY26. Observations for MY27 are only available until Ls = 82◦ , butthis period displays slightly lower dust optical depth during earlysummer compared to prior years. The spatial and temporal distri-bution of dust optical depth events indicates that these increasesare likely due to cap edge storms, induced by thermal gradients re-sulting from the expansion and regression of the polar cap (Jameset al., 2005). Fig. 7b shows the spatial distribution of dust opac-ity for regular intervals in the vicinity of the South pole, for thecomplete martian seasonal cycle.

Prominent repeating water ice clouds were observed between25◦ and 45◦ S latitude during late autumn to early winter of eachyear (Ls = 70◦ to 140◦). Water ice activity was stronger polewardof 60◦ S latitude, between Ls = 350◦ and 171◦ in MY25 than inother years. A reduction of around 0.2 in water ice optical depthoccurred at Ls = 100◦–169◦ in MY26 and was the only year to ex-hibit this behavior. This reduction in water ice was coincident witha relatively large increase (0.5–0.7) in dust optical depth.

Also of note are the possible detections of water ice at Ls =80◦–230◦ in MY24–26, which extended from 25◦ S to 42◦ S atpeak activity (see Fig. 5). Despite the gap due to poor thermal con-trast, this feature appears to be continuous between the results ofthe new algorithm and the lower latitude survey of ice aerosols by(Smith, 2004).

Temperature increases were more widespread in the vicinity ofthe South pole. A better correlation between temperature and dustoptical depth was detected around the South pole compared tothe North. Fig. 5 shows the greater coincidence in the size andtiming of southern temperature increases. Although, temperaturecontrasts between regions/times remained the same as those inthe North at ≈10–20 K.

4.3. Annular bands

The most striking difference in dust activity in the vicinity ofthe poles is the formation of “clear annular bands” from 45◦ Nto 60◦ N, Ls = 220◦–310◦ as seen in Figs. 7a and 7b. These clearbands tend to form during the northern autumn–winter periodfrom around Ls = 250◦–310◦ and follow much the same morphol-ogy in each year.

Clear bands seem to span a significant temperature gradient.For the specific case of MY24, the edge of the band lies at 61◦ N

Table 2Maximum latitudinal elongation of northern clear annular bands and duration ofeach band.

Mars Year Maximum elongation south (degrees) Duration (Ls)

MY24 16◦ 122◦MY25 13◦ 107◦MY26 14◦ 108◦

and is subject to surface temperatures of around 150 K, whileequivalent temperatures adjacent to the bottom edge of the for-mation at 46◦ N are around 170–180 K. Although similar extremetemperature gradients can be observed in the vicinity of the Southpole, dust opacity banding appears to occur only around the Northpole and was not observed in the South. The annular bands inMY25 demonstrated a reduced latitudinal component comparedwith MY24 and MY26. This is likely due to the increased levels ofairborne dust injected into atmosphere by the 2001a dust storm.

Table 2 shows the maximum latitudinal and temporal extent ofthese bands for each surveyed Mars Year. Alterations in the mor-phology of the annular banding may be due to changes in thethermal response of the surface or the result of local circulationchanges between years. Any thermal changes present in the re-gion have a minimal effect on the polar cap itself as no significantchange is detected in regression rate (Benson and James, 2005).

5. Conclusions

• Modifications made to adapt the standard TES retrieval algo-rithm (Smith et al., 2000a) to account for the spectral slopepresent in cold surface spectra are valid. The modified algo-rithm is capable of retrieving aerosol optical depth over sur-faces with temperatures below 210 K.

• Reliable observations of dust optical depth have been obtainedover both martian poles from late MY24 to early MY27. Theseobservations contribute to the overall completeness of the TESdataset.

• The greatest observed dust activity each year takes place pole-ward of 40◦ N latitude from Ls = 170◦ to 30◦ and poleward of50◦ S from Ls = 340◦ to 180◦ .

• Dust optical depths near the South pole climbed sharply at thebeginning of each year likely due to thermal gradients causedby the rapid regression of the polar cap over this period.

• There is a strong correlation between temperature contrastand high dust optical depth. This effect is most pronouncedduring the expansion and regression of the polar caps fromLs = 140◦ to 270◦ and Ls = 340◦ to 90◦ as dust storms are in-duced by the temperature gradients present along the polarcap edge. Surface temperatures are elevated in the vicinity ofa dust event.

• The widespread increases of dust optical depth seen aroundthe North pole in MY26 may be due to increased amounts ofsurface dust distributed by the 2001a dust storm of the previ-ous year. The thermal properties of the surface may also haveundergone modification due to increased dust deposition inMY25. This may be responsible for the ‘patchy’ surface warm-ing seen for MY26 in Fig. 5.

• Consistently repeating annular bands of low dust and ice op-tical depth relative to the surrounding areas formed between45◦ N and 60◦ N latitude from mid-autumn (Ls = 220◦) tomid-winter (Ls = 310◦) in each Mars Year.

• The annular low optical depth bands around the North polecould result from the formation of a “flushing” dust storm.These may be generated by frontal dust storms which resem-ble standard baroclinic fronts that grow into regional scalestorms. North polar cap edge storms have been confirmed byMOC from Ls = 200◦ to 240◦ (north autumn/south spring) and

128 D. Horne, M.D. Smith / Icarus 200 (2009) 118–128

Ls = 300◦ to 340◦ (north winter/south summer) (Cantor et al.,2001; Wang et al., 2003), coincident with our observationof low optical depth bands. General Circulation Models havealso successfully demonstrated the effect of a flushing storm(Newman et al., 2002), indicating that dust may be raised bytemperature induced cap edge winds and channeled south-wards away from the pole as it is picked up by the globalHadley cycle.

• The annular bands were visibly affected in physical size andlongevity by the effects of the 2001a dust storm.

• The presence of annular bands in the vicinity of the Northpolar cap and their absence over the South pole may be in-dicative of differences in hemispherical circulation patterns orthe individual thermal response of the North and South polarsurfaces.

• Interannual changes were observed in the morphology of theannular banding. These changes were due to differences indust activity in the upper and lower bands. While changes inMY25 may be a result of the 2001a dust storm, no such eventtook place in MY26. It is possible that interannual modificationin the thermal response of the surface may be responsible forinducing dust activity.

• The annular banding pattern over the North pole appears tocorrelate with strong cross latitude thermal contrasts; this maybe indicative of local and global circulation effects.

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