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Water and the Martian W cloud
ICARlJS,
18,497~501
(1973)
Water
and
the
Martian
W Cloud
S. J. PEALE Department Santa Received
July
The diurnal brightening Mariners 6 and 7 is likely required in the cloud and brightening is estimated t,o local source with t,his daily daily uplift, and saturation tions.
of Physics, Barbara,
University California
30, 1972;
revised
of the W due to the as vapor in bp roughly production of existing
Leovy et al. (197 1) have described several equatorial regions on the surface of Mars near 1lO”W longitude whose surface brightness shows a diurnal variation with a maximum in the late afternoon. These regions compose parts of the W cloud defined by earth-based observations. The afternoon brightenings are evident on the sequences of Mariner 6 and 7 photographs 6F37-6F39 and 7F77-7F79 which show an increase in the relative reflectance of as much as 0.06. Relative reflectance is clefined by R = n-H/F,, where B is the observed surface brightness and F, is the incident flux densitS of solar radiation. R is thus the ra,tio of the observed surface brightness to that of a white Lambert surface illuminat,ed by F,. The unique character of these brightenings, the unlikeliness of their being CO, condensations, and their correlation with a warming and subsequently cooling surface leads to the possibility that’ they may represent a’ condensed phase of a daily rxchange of‘ water vapor between t.hc ground and the atmosphere. The rapid depletion of any permafrost source of water necessitates a deeper source mnintained in a thermally active region it significant amounts of water are to be outgassed by the surface (Leovy et al.: Q 1973 by Academic Press, Inc. of reproduction in any form reserved.
October
24,
1972
cloud region of Mars during the flyby of formation of water ice clouds. The water the atmosphere to produce the observed bet,ween 0.1 and 1 Opm. Eit’her a qualified over the W cloud area or a more probable atmosphere are compatible with observa-
I. INTRODUCTION
Copyright All rights
of California, 93106
497
1971). It is significant that the W cloud forms over a high volcanic plateau characterized by the three large colinear craters southeast of Nix Olympica. The necessary thermal conditions for the current outgassing of water vapor may persist in this region. In Section II we use the work of van de Hulst (1948) to estimate a lower bound on the amount of atmospheric water in the form of ice particles which could cause the observed brightness increase. This section ends with a reservation about excluding CO, condens&ion as the cause of the brightness increase. With water the assumed cause of the brightening phenomena, the total atmospheric water content during the brightest phase is estimated in Section III as a function of cloud altitude. These results are discussed in Section IV, where a possible active source of outgassing water vapor is qualified but concluded to be unnecessary for the condensation of water ice clouds. Sufficient water already exists in the atmosphere and the high local topography can most likely induce the necessary diurnal elevation and cooling Do saturation. On the assumpt,ion that, the brightenrtl regions are due to clouds in the atmosphere we must determine the optical depth of the cloud cover which is necessary to
498
PEALE
produce the observed increase in relative reflectance. The problem of a multiply scattering atmosphere above a diffusely reflecting planetary surface has been investigated by van de Hulst (1948) under the assumptions that the surface reflects like a Lambert surface of albedo b and that the atmospheric particles scatter isotropically with albedo a. Use of these assumptions minimizes the optical depth required for a given increase in relative reflectance (since real cloud particles will be forward scattering) and gives a lower bound on the necessary amount of aerosol. Leovy et al. (1971) found the relative reflectance of the W cloud region to increase on two separate days from approximately 0.16 to 0.22 with maximum occurring as the region approaches the limb at about ~:OOPM local Martian time. We shall assume no contribution to the surface brightness from atmospheric aerosols prior to the brightening and neglect any limb darkening effects due to the atmosphere. The Mariner photograph 7F79 corresponding to maximum brightness of the W cloud was taken with a solar phase angle of 23” when the central longitude was 195’W. The central part of the W cloud was located 83” further east on the equator. The necessary parameters used for determining the optical depth for a relative reflectance R of 0.22 in the brightened region follow from the above data and are given in Eq. (1) where p is the cosine of the the downward vertical angle between and the direction to the observer, and p’ is bhe cosine of the angle between the downward vertical and the incident solar ratiation. The aerosol albedo has been ‘1
R = 0.22 corresponds to T= 0.02. From this value of r, we can estimate the area1 mass density of the cloud particles responsible for the .brightness increase. The optical depth can be expressed 7 = nax,
where n is the mean particle density, u the mean cross section per particle: and x the vertical depth of the region containing aerosol. The mass in a vertical column per unit area of surface is thus pa = $TpT:
(‘1 arbitrari1.v set to unity H-hi& minimizes the optical depth 7. The approximation valid for a~ Q 1 was used to calculate H for several values of 7 witoh the result, that
(3)
where r is the mean particle radius and p is the (volume) mass density. If the particles are water ice and we minimize the mass by choosing the rather small value of r = lo-” cm, the area1 mass density of the cloud particles is about 1.7 x 10-6g/cm2 = 2.7 X 10P2pm of precipitable water. The assumptions above are always weighted toward minimizing the amount of necessary cloud mass, so the precipitable water in the form of ice particles should be considered an absolute lower bound. The mass densities of other cloud materials such as solid CO, or dust can also be estimated by the above procedure. However, the 33-km altitude of the frost point of CO, (Hogan et al., 1972) would have made CO, clouds visible on the limb, which the W cloud was not (Leovy et al., 1971). On the other hand, the altitude of the frost point of CO, could be considerably lower (22- to 25-km altitude) if the temperature gradient were increased to the adiabatic one by some type of forced convection (Gierasch, 197 1). This possibility makes the limb observation perhaps less definitive in excluding CO, clouds, but in any case we shall concentrate on testing H,O as the source of the diurnal brightening. III.
b = 0.16.
(2)
NE(T~~ARY
\VATER
VAPOR
In order to estimate the total water rtlrjuired to give the diurnal brightening, we shall assume saturation at the level of a thin cloud (say .NOm thick). a constant mixing ratio of H,O to t,he ground, and
WATER
AND
THE
MARTIAN
no water vapor above the cloud level. Since the W cloud region is on the Martian equator and the Mariner 6 entry was also equatorial, the thermal profile determined by the Mariner 6 entry is an appropriate reference. Although the measured temperature gradient at the Mariner 6 entry point was subadiabatic at 3”/km (Kliore et al., 1969; Hogen et al., 1972) the formation of clouds implies convective lift of atmosphere containing water vapor from lower elevations and the establishment of an adiabatic lapse rate. We determine the atmospheric water content for two assumed atmospheric temperatures at an (estimated) 8 km mean altitude of the volcanic plateau under the VI’ cloud. All altitudes refer to the Mariner 6 occultation as zero elevation at an equatorial radius of 3394km. One of the assumed base t#emperatures of 225°K corresponds to the atmospheric temperature at 8km as determined by the Mariner 6 entry. The resulting thermal structure with an adiabatic lapse rate would imply little or no thermal coupling of the atmosphere to the ground and might result from an upward deflection of horizontal winds with adiabatic cooling. The second base temperature of 246°K corresponds to the atmospheric temperature of 184°K at 20 km at the Mariner 6 entry point with an adiabatic lapse rate of 5.16”K/km determining the temperature of the plateau level of 8 km. This thermal structure would imply strong thermal coupling to the ground with convective equilibrium being established up to the cloud level. If an adiabatic gradient can be established at all, the base temperature, which is not necessarily the ground temperature (Gierasch and Goody, 1971; Hogan et al., 1972), is likely to fall within these extremes. Combining the adiabatic relation between temperature and pressure with the equation of hydrostatic equilibrium yields t,he above adiabatic lapse rate of 5.16”K/km for a CO, a,tmosphere wibh the ratio of specific heats y = 1.35. This allows us to write the atmospheric temperature at altitude h as T = T, - 5.16 x lo-‘(h
- h,).
t-11
W
499
CLOUD
where T, is altitude h, and in centimeters. of the adiabatic ture T is given
the base temperature at where hand h, are expressed The total number density atmosphere at a temperaby
(5) where n,- is the density at temperature T,. The number density of water molecules is then (6) where the subscript c denotes “cloud” and where a constant mixing ratio equal to that at the cloud level has been assumed. The density nwc is the saturation density of water vapor at the cloud level and n, follows from Eq. (5) with T = T,. From the Clausius-Clapeyron relation the vapor pressure of ice at temperature I’ is *, = Ke-M’lNkT > (7) where K is a constant, 1 is the latent heat of sublimation, N is Avogadro’s number, k is Boltzman’s constant, and .iW is the molecular weight of water. We choose I= 676cal/g the value at the triple point and evaluate the con tant K from the “f, known triple-point temperature and pressure. Thus 3 . yrj . * y ]~13e-6110/T~ rzwc
=
---- crnm3
--
(8)
kTc
y’, -. r,. 16 x I()-’ (h, . lb,,)‘ l.‘(Y-~1) -.- ~~~,~< --~--.-~~i) t lc
(9)
\vherr Eq. (4) has been used. The total water vapor in a vertical column of one ‘i( Iuarc cent,imet,er cross section is thus
500
PEALE
The total water which is necessary to produce the diurnal brightening is shown in Fig. 1 as a function.of altitude for the two base temperatures discussed above. The horizontal line is the water in the form of ice particles and is seen to be a small fraction of the total except at the higher cloud altitudes. This effectively places a bound on any water cloud altitude of about 18-23 km since it is almost certainly true that most of the water in the entire depth of atmosphere will not be in the condensed phase. This range of upper bounds on the cloud altitude is consistent. with the upper bound of 25 km from the observations (Leovy et al., 1971). If one accounts for the forward scattering of the ice particles, the resulting increase in optical depth above the lower bound of 0.02 would force the clouds to still lower altitudes to maintain a large dominance of the vapor phase. IV.
DISC~BSION
From Fig. 1, we see that something between 0.1 and 10~ of precipitable water is all that is necessary to yield the observed diurnal brightening of the W cloud region, We estimate the area of the brightened regions from Leovy et al. ( 197 1) to be about IO5 km2. If the above water is assumed to be produced daily from vents on the surface, the given area implies a source of 10’0~~10’2g/day or about
3.3 x 1033-3.3 x 1O3s molecules/day. The observed loss rate of hydrogen from the planet is 2 x lo8 atoms (cm2 day) (Barth et al., 1972) which implies a loss of 103’ molecules of water per day over the entire planet if the source of the escaping hydrogen is photodissociated water (Hunten and McElroy, 1971). The hypothesized source of H,O exceeds the escape rate by several orders of magnitude and consistency would require that such a source be active only at rare intervals or that an efficient transport mechanism allows the excess to be fixed in the permanent north polar cap. In the situation where the volcanic plateau under the W cloud is the only active source of water on the entire planet, approximately 10-1000 days would be required to yield 1 pm of precipitable water distributed over the entire planet corresponding to the O.l- to lopm/day source over the W cloud region. This time scale may be sufficient to allow transport of water to a polar sink while preventing an accumulation in the atmosphere above that which is observed. An active source of water may thus be sufficiently qualified to account for the W cloud. An alternative and perhaps more probable explanation of the diurnal brightening follows from the small amount of water required (Fig. l), and from the topography of the W cloud region. The higher estimate of 1Opm is comparable to the
CLOUD CONTENT I@-
I$ 6
12
16
20
24
cloud
ultitude.
WATER
AND
THE
amount of water often observed in the Martian atmosphere (Owen and Mason, 1969; Back et al., 1970). If this amount of water is already in the atmosphere, there is obviously no need for a local source with this diurnal production for the W cloud. Adiabatic cooling of the existing water-laden atmosphere to the saturation temperature would suffice. Likely mechanisms for such cooling are the elevation of diurnally enhanced trade winds as they rise over the high plateau or the elevation of low-lying air masses by a local, thermally induced circulation. The latter may follow if there is sufficient thermal coupling of the atmosphere to the ground generating large horizontal temperature gradients in regions of large elevation differences. Such a flow has been described by Gierasch and Sagan (197 1) and would have the necessary upslope direction during the day. In fact, the correlation of the Martian white clouds with high regions on the surface led Sagan et al. (1971) to propose such adiabatic cooling of condensibles as a probable mechanism for the cloud formation. If we confine 1Opm of water in the lower 2km of atmosphere above zero elevation, using a pressure of 5.8 mb (at l-km elevation) leads to a mixing ratio of about 10p3. A slight modification of Eq. (9) then yields the elevation of about 9km for saturation if this air mass were lifted and cooled adiabatically. (The temperature at zero elevation was assumed to be the Mariner 6 entry value of 250°K.) A plateau altitude exceeding 8 km is entirely adequate for saturation of lower-lying air masses forced to pass over it. The fact that any air mass will rise to somewhat higher altitudes than t,he maximum ground level over which it is forced to pass allows lower mixing ratios than the 10e3 used above to still produce the necessary clouds. Wat,er icar clouds are the likely cause of thr diurnal brightening of the W cloud Xthough t,he t’otal necessary region. amount of wat’ei involved in t’he phenomena is small enough to perhaps qualify a locajl active source, the small value implies that cooling and saturation of the existing
MARTIAN
W
CLOUD
501
atmosphere as it rises over the high plateau is an adequate and more likely explanation. Mapping of the region with an infrared spectrometer with high spatial resolution would detect any horizontal gradients in atmospheric water vapor content and may comprise a definitive test for a local source of water. ACKNOWLEDGMENT This
research
is funded by of Space NGR 05-010-062
Program Office, Office under
Grant
the Planetoloa Science, NASA,
REFERENCES
B.IRCH, C. A..
STEWART, A. I., HORD, C. TV., AND LANE, A. L. (1972). Mariner 9 ultraviolet experiment : Mars airglow spectroscopy and variations in Lyman a. Icarus, in press. BARKER, E. S., SCHORN, R. A., WOSZCZYK, A., TULL, R. G., AND LITTLE, S. J . (1970). Mars: Detection of atmospheric water vapor during the southern hemisphere Spring and Summer season. Science 170, 1308. GIERASCH, P. J., AND GOODY, R. M. (1968). A study of the thermal and dynamical structure of the Martian lower atmosphere. Planet. Space Sci. 16, 615. GIERASCH, P., AND SAGEN, C. (1971). A preliminary assessment of Martian wind regimes. Icarus 14, 372. GIERASCH, P. (1971). Dissipationin atmospheres: The thermal structure of the Martian lower atmosphere with and wighout viscous dissipation. J. Atmos. Sci. 28, 315. HOGAN, J. S., STEWART, R. W., AND RASOOL. S. I. (1972), Radio occultation measurements of the Mars atmosphere with Mariners 6 and 7. Radio Sci., in press. HULST, H. C. VBN DE (1948). Scattering in a planetary atmosphere. A&on. J. 107, 220. HUNTEN, ID. M., AND MCELROY, M. B. (1970). Production a,nd escape of hydrogen on Mars. J. Geophys. Res. 75, 5989. KIJORE, A., FJELDBO. G., AND SEIDEI~, B. 1~. (1969). Mariners 6 and 7 : Radio occultation measurements of the atmosphere of Mars. Science 166, 1393. I,EOVV, C?., SRIITH, B. A.. YOIJFX, A. 'I'.. AXD T,F:IOHTOS. X. R. (1971). $fariner Mars 1969: ;Itmosphc:ric results. ,/. C:eophys. Kes. 76, 297. I )WIW. T.. AXD ,\l.wox. K. P. ( 1969). Mars: \i-attli, \-apnr ix, it > iitmonphcr~ .S’risnc~ 165. 893. SAGAN. C'.. VEVERKA. ,I., AXI) GIERASCH, P. (197 1). Observat,ional consequences of Martian \vind regimes. .Icar7ts 15, 253.-278.
18,497~501
(1973)
Water
and
the
Martian
W Cloud
S. J. PEALE Department Santa Received
July
The diurnal brightening Mariners 6 and 7 is likely required in the cloud and brightening is estimated t,o local source with t,his daily daily uplift, and saturation tions.
of Physics, Barbara,
University California
30, 1972;
revised
of the W due to the as vapor in bp roughly production of existing
Leovy et al. (197 1) have described several equatorial regions on the surface of Mars near 1lO”W longitude whose surface brightness shows a diurnal variation with a maximum in the late afternoon. These regions compose parts of the W cloud defined by earth-based observations. The afternoon brightenings are evident on the sequences of Mariner 6 and 7 photographs 6F37-6F39 and 7F77-7F79 which show an increase in the relative reflectance of as much as 0.06. Relative reflectance is clefined by R = n-H/F,, where B is the observed surface brightness and F, is the incident flux densitS of solar radiation. R is thus the ra,tio of the observed surface brightness to that of a white Lambert surface illuminat,ed by F,. The unique character of these brightenings, the unlikeliness of their being CO, condensations, and their correlation with a warming and subsequently cooling surface leads to the possibility that’ they may represent a’ condensed phase of a daily rxchange of‘ water vapor between t.hc ground and the atmosphere. The rapid depletion of any permafrost source of water necessitates a deeper source mnintained in a thermally active region it significant amounts of water are to be outgassed by the surface (Leovy et al.: Q 1973 by Academic Press, Inc. of reproduction in any form reserved.
October
24,
1972
cloud region of Mars during the flyby of formation of water ice clouds. The water the atmosphere to produce the observed bet,ween 0.1 and 1 Opm. Eit’her a qualified over the W cloud area or a more probable atmosphere are compatible with observa-
I. INTRODUCTION
Copyright All rights
of California, 93106
497
1971). It is significant that the W cloud forms over a high volcanic plateau characterized by the three large colinear craters southeast of Nix Olympica. The necessary thermal conditions for the current outgassing of water vapor may persist in this region. In Section II we use the work of van de Hulst (1948) to estimate a lower bound on the amount of atmospheric water in the form of ice particles which could cause the observed brightness increase. This section ends with a reservation about excluding CO, condens&ion as the cause of the brightness increase. With water the assumed cause of the brightening phenomena, the total atmospheric water content during the brightest phase is estimated in Section III as a function of cloud altitude. These results are discussed in Section IV, where a possible active source of outgassing water vapor is qualified but concluded to be unnecessary for the condensation of water ice clouds. Sufficient water already exists in the atmosphere and the high local topography can most likely induce the necessary diurnal elevation and cooling Do saturation. On the assumpt,ion that, the brightenrtl regions are due to clouds in the atmosphere we must determine the optical depth of the cloud cover which is necessary to
498
PEALE
produce the observed increase in relative reflectance. The problem of a multiply scattering atmosphere above a diffusely reflecting planetary surface has been investigated by van de Hulst (1948) under the assumptions that the surface reflects like a Lambert surface of albedo b and that the atmospheric particles scatter isotropically with albedo a. Use of these assumptions minimizes the optical depth required for a given increase in relative reflectance (since real cloud particles will be forward scattering) and gives a lower bound on the necessary amount of aerosol. Leovy et al. (1971) found the relative reflectance of the W cloud region to increase on two separate days from approximately 0.16 to 0.22 with maximum occurring as the region approaches the limb at about ~:OOPM local Martian time. We shall assume no contribution to the surface brightness from atmospheric aerosols prior to the brightening and neglect any limb darkening effects due to the atmosphere. The Mariner photograph 7F79 corresponding to maximum brightness of the W cloud was taken with a solar phase angle of 23” when the central longitude was 195’W. The central part of the W cloud was located 83” further east on the equator. The necessary parameters used for determining the optical depth for a relative reflectance R of 0.22 in the brightened region follow from the above data and are given in Eq. (1) where p is the cosine of the the downward vertical angle between and the direction to the observer, and p’ is bhe cosine of the angle between the downward vertical and the incident solar ratiation. The aerosol albedo has been ‘1
R = 0.22 corresponds to T= 0.02. From this value of r, we can estimate the area1 mass density of the cloud particles responsible for the .brightness increase. The optical depth can be expressed 7 = nax,
where n is the mean particle density, u the mean cross section per particle: and x the vertical depth of the region containing aerosol. The mass in a vertical column per unit area of surface is thus pa = $TpT:
(‘1 arbitrari1.v set to unity H-hi& minimizes the optical depth 7. The approximation valid for a~ Q 1 was used to calculate H for several values of 7 witoh the result, that
(3)
where r is the mean particle radius and p is the (volume) mass density. If the particles are water ice and we minimize the mass by choosing the rather small value of r = lo-” cm, the area1 mass density of the cloud particles is about 1.7 x 10-6g/cm2 = 2.7 X 10P2pm of precipitable water. The assumptions above are always weighted toward minimizing the amount of necessary cloud mass, so the precipitable water in the form of ice particles should be considered an absolute lower bound. The mass densities of other cloud materials such as solid CO, or dust can also be estimated by the above procedure. However, the 33-km altitude of the frost point of CO, (Hogan et al., 1972) would have made CO, clouds visible on the limb, which the W cloud was not (Leovy et al., 1971). On the other hand, the altitude of the frost point of CO, could be considerably lower (22- to 25-km altitude) if the temperature gradient were increased to the adiabatic one by some type of forced convection (Gierasch, 197 1). This possibility makes the limb observation perhaps less definitive in excluding CO, clouds, but in any case we shall concentrate on testing H,O as the source of the diurnal brightening. III.
b = 0.16.
(2)
NE(T~~ARY
\VATER
VAPOR
In order to estimate the total water rtlrjuired to give the diurnal brightening, we shall assume saturation at the level of a thin cloud (say .NOm thick). a constant mixing ratio of H,O to t,he ground, and
WATER
AND
THE
MARTIAN
no water vapor above the cloud level. Since the W cloud region is on the Martian equator and the Mariner 6 entry was also equatorial, the thermal profile determined by the Mariner 6 entry is an appropriate reference. Although the measured temperature gradient at the Mariner 6 entry point was subadiabatic at 3”/km (Kliore et al., 1969; Hogen et al., 1972) the formation of clouds implies convective lift of atmosphere containing water vapor from lower elevations and the establishment of an adiabatic lapse rate. We determine the atmospheric water content for two assumed atmospheric temperatures at an (estimated) 8 km mean altitude of the volcanic plateau under the VI’ cloud. All altitudes refer to the Mariner 6 occultation as zero elevation at an equatorial radius of 3394km. One of the assumed base t#emperatures of 225°K corresponds to the atmospheric temperature at 8km as determined by the Mariner 6 entry. The resulting thermal structure with an adiabatic lapse rate would imply little or no thermal coupling of the atmosphere to the ground and might result from an upward deflection of horizontal winds with adiabatic cooling. The second base temperature of 246°K corresponds to the atmospheric temperature of 184°K at 20 km at the Mariner 6 entry point with an adiabatic lapse rate of 5.16”K/km determining the temperature of the plateau level of 8 km. This thermal structure would imply strong thermal coupling to the ground with convective equilibrium being established up to the cloud level. If an adiabatic gradient can be established at all, the base temperature, which is not necessarily the ground temperature (Gierasch and Goody, 1971; Hogan et al., 1972), is likely to fall within these extremes. Combining the adiabatic relation between temperature and pressure with the equation of hydrostatic equilibrium yields t,he above adiabatic lapse rate of 5.16”K/km for a CO, a,tmosphere wibh the ratio of specific heats y = 1.35. This allows us to write the atmospheric temperature at altitude h as T = T, - 5.16 x lo-‘(h
- h,).
t-11
W
499
CLOUD
where T, is altitude h, and in centimeters. of the adiabatic ture T is given
the base temperature at where hand h, are expressed The total number density atmosphere at a temperaby
(5) where n,- is the density at temperature T,. The number density of water molecules is then (6) where the subscript c denotes “cloud” and where a constant mixing ratio equal to that at the cloud level has been assumed. The density nwc is the saturation density of water vapor at the cloud level and n, follows from Eq. (5) with T = T,. From the Clausius-Clapeyron relation the vapor pressure of ice at temperature I’ is *, = Ke-M’lNkT > (7) where K is a constant, 1 is the latent heat of sublimation, N is Avogadro’s number, k is Boltzman’s constant, and .iW is the molecular weight of water. We choose I= 676cal/g the value at the triple point and evaluate the con tant K from the “f, known triple-point temperature and pressure. Thus 3 . yrj . * y ]~13e-6110/T~ rzwc
=
---- crnm3
--
(8)
kTc
y’, -. r,. 16 x I()-’ (h, . lb,,)‘ l.‘(Y-~1) -.- ~~~,~< --~--.-~~i) t lc
(9)
\vherr Eq. (4) has been used. The total water vapor in a vertical column of one ‘i( Iuarc cent,imet,er cross section is thus
500
PEALE
The total water which is necessary to produce the diurnal brightening is shown in Fig. 1 as a function.of altitude for the two base temperatures discussed above. The horizontal line is the water in the form of ice particles and is seen to be a small fraction of the total except at the higher cloud altitudes. This effectively places a bound on any water cloud altitude of about 18-23 km since it is almost certainly true that most of the water in the entire depth of atmosphere will not be in the condensed phase. This range of upper bounds on the cloud altitude is consistent. with the upper bound of 25 km from the observations (Leovy et al., 1971). If one accounts for the forward scattering of the ice particles, the resulting increase in optical depth above the lower bound of 0.02 would force the clouds to still lower altitudes to maintain a large dominance of the vapor phase. IV.
DISC~BSION
From Fig. 1, we see that something between 0.1 and 10~ of precipitable water is all that is necessary to yield the observed diurnal brightening of the W cloud region, We estimate the area of the brightened regions from Leovy et al. ( 197 1) to be about IO5 km2. If the above water is assumed to be produced daily from vents on the surface, the given area implies a source of 10’0~~10’2g/day or about
3.3 x 1033-3.3 x 1O3s molecules/day. The observed loss rate of hydrogen from the planet is 2 x lo8 atoms (cm2 day) (Barth et al., 1972) which implies a loss of 103’ molecules of water per day over the entire planet if the source of the escaping hydrogen is photodissociated water (Hunten and McElroy, 1971). The hypothesized source of H,O exceeds the escape rate by several orders of magnitude and consistency would require that such a source be active only at rare intervals or that an efficient transport mechanism allows the excess to be fixed in the permanent north polar cap. In the situation where the volcanic plateau under the W cloud is the only active source of water on the entire planet, approximately 10-1000 days would be required to yield 1 pm of precipitable water distributed over the entire planet corresponding to the O.l- to lopm/day source over the W cloud region. This time scale may be sufficient to allow transport of water to a polar sink while preventing an accumulation in the atmosphere above that which is observed. An active source of water may thus be sufficiently qualified to account for the W cloud. An alternative and perhaps more probable explanation of the diurnal brightening follows from the small amount of water required (Fig. l), and from the topography of the W cloud region. The higher estimate of 1Opm is comparable to the
CLOUD CONTENT I@-
I$ 6
12
16
20
24
cloud
ultitude.
WATER
AND
THE
amount of water often observed in the Martian atmosphere (Owen and Mason, 1969; Back et al., 1970). If this amount of water is already in the atmosphere, there is obviously no need for a local source with this diurnal production for the W cloud. Adiabatic cooling of the existing water-laden atmosphere to the saturation temperature would suffice. Likely mechanisms for such cooling are the elevation of diurnally enhanced trade winds as they rise over the high plateau or the elevation of low-lying air masses by a local, thermally induced circulation. The latter may follow if there is sufficient thermal coupling of the atmosphere to the ground generating large horizontal temperature gradients in regions of large elevation differences. Such a flow has been described by Gierasch and Sagan (197 1) and would have the necessary upslope direction during the day. In fact, the correlation of the Martian white clouds with high regions on the surface led Sagan et al. (1971) to propose such adiabatic cooling of condensibles as a probable mechanism for the cloud formation. If we confine 1Opm of water in the lower 2km of atmosphere above zero elevation, using a pressure of 5.8 mb (at l-km elevation) leads to a mixing ratio of about 10p3. A slight modification of Eq. (9) then yields the elevation of about 9km for saturation if this air mass were lifted and cooled adiabatically. (The temperature at zero elevation was assumed to be the Mariner 6 entry value of 250°K.) A plateau altitude exceeding 8 km is entirely adequate for saturation of lower-lying air masses forced to pass over it. The fact that any air mass will rise to somewhat higher altitudes than t,he maximum ground level over which it is forced to pass allows lower mixing ratios than the 10e3 used above to still produce the necessary clouds. Wat,er icar clouds are the likely cause of thr diurnal brightening of the W cloud Xthough t,he t’otal necessary region. amount of wat’ei involved in t’he phenomena is small enough to perhaps qualify a locajl active source, the small value implies that cooling and saturation of the existing
MARTIAN
W
CLOUD
501
atmosphere as it rises over the high plateau is an adequate and more likely explanation. Mapping of the region with an infrared spectrometer with high spatial resolution would detect any horizontal gradients in atmospheric water vapor content and may comprise a definitive test for a local source of water. ACKNOWLEDGMENT This
research
is funded by of Space NGR 05-010-062
Program Office, Office under
Grant
the Planetoloa Science, NASA,
REFERENCES
B.IRCH, C. A..
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