JOURNAL OF GEOPHYSICAL RESEARCH, VOL. 104, NO. E4, PAGES 9043-9053, APRIL 25, 1999

Cloud formation under Mars Pathfinder conditions

Anthony Colaprete and Owen B. Toon Laboratory for Atmospheric and Space Physics, University of Colorado, Boulder

Julio A. Magalh•es Space Sciences Division, NASA Ames Research Center, Moffett Field

Abstract. Temperature data taken during the Mars Pathfinder entry are applied to a time dependent microphysical model of cloud formation. Our goal is to reproduce the characteristics of clouds observed at the Pathfinder site. We assume that the total water column abundance is 11

percipitable microns. We also assume that the atmospheric dust initially has a lognormal size distribution with a modal radius of 0.6 gm, reff = 1.5 gm, and an optical depth of q: = 0.5. Three temperature profiles are used in the model calculations. The first is the Pathfinder entry observation. The second is this same profile but with the inversion at 10 km removed. The third simulation uses the temperature profiles calculated by the Ames Mars general circulation model for Mars Pathfinder conditions. In all cases clouds form at altitudes of 35-45 km. In addition lower thicker clouds form within the 10 km inversion of the unaltered profile, while peak cloud formation occurred near local noon at 20 km for the GCM calculated l•rofiles. Given a combination of all three temperature scenarios, cloud optical depths, particle sizes, and colors are consistent with IMP observations.

1. Introduction

Clouds are an important element of the Martian climate system because of their ability to redistribute water vapor [James, 1990; Kahn, 1990; Jakosky 1983; Michelangeli et al., 1993; Clancy et al., 1996]. As water vapor condenses, clouds of ice particle form. These ice crystals fall until they either evaporate or reach the ground, redistributing water vapor from higher to lower altitudes. This redistribution may have an effect on the division of mass between the north and south polar caps and on the observed asymmetry between the northern and southern hemisphere's atmospheric water vapor abundances [James, 1990; Clancy et al., 1996]. By driving the water vapor to lower altitudes, Martian clouds facilitate the exchange of water vapor with the Martian regolith [Kahn, 1990], possibly explaining the striking decline in water abundance observed during the late northern spring [Jaksoky, 1983]. Since it is likely that the condensation of ice particles is occurring on suspended dust within the atmosphere, cloud formation will also modify the amount and location of dust as it "sweeps" it from the sky. This modification in the distribution of dust could effect the formation of local dust

storms, contribute to the inter-annual variability of global dust storms [Clancy et al., 1996] and be responsible for variations in the polar ice cap laminae [Pollack et al., 1981].

While observations of clouds and hazes date back to Mariner 9

and Viking in the 1970s, the variability of these clouds, both temporally and spatially, had not been fully appreciated until recently. The Imager for Mars Pathfinder (IMP) has shown a

Copyfight 1999 by the American •3eophysical Union.

Paper number 1998JE900018. 0148-0227/99/1998JE900018509.00

highly variable and dynamic cloud system that may be coupled with both diurnal and seasonal fluctuations in atmospheric temperature. These observations, combined with the temperature profile measured by the Mars Pathfinder Atmospheric Structures Investigation/Meteorology (ASI/MET) package, allow improved modeling of Martian cloud formation and evolution. Using these data we explore two questions: Where and when do clouds form under Mars Pathfinder conditions? What are the physical and optical properties of these clouds? Our model results are consistent with the observations of optical depth and cloud color made by IMP.

2. Observations

ASI/MET [Magalhaes et al., this issue] measured the Martian atmospheric temperature profile during descent on July 4, 1997 (see Figure 1). While in some respects the Pathfinder temperature profile is similar to the temperature profiles measured by the two Viking entry probes, several differences exist. Wave features above about 20 km are evident, as they are in the Viking data sets; however, the coldest Pathfinder temperatures are 40 K lower than the Viking temperatures. Another feature of the Pathfinder data is an inversion at about 10

km which is not present in either of the Viking entry profiles. This inversion leads to the formation of a low-altitude cloud deck

in our model.

IMP made numerous observations of clouds in the predawn hours and also after sunset. These clouds varied in spatial extent from near complete overcast to singular wisps. The color of the clouds alternated from near white to pink and, at some times, blue. The shape of the clouds ranged from scattered and fluffy to more confined and lens-like.

9043

9044 COLAPRETE ET AL.: CLOUD FORMATION AT PATHFINDER

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• 20

100 120 140 160 180 200 220

Tempe•anu•e (K)

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,

. . i . . . i . . . i . . . i . . . i . . •,. i

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Figure 1. •e three temperature profiles used in this work. (a) •e solid curve is the measured temperature profile by ASI•ET (see text). •e dotted and dashed curves are the average and maximum temperatures, respectively, utilized by the model. (b) The solid curve is the ASI/MET data with the lowest inversion

removed. •e dotted and dashed curves are the same as in Figure la above the inversion. (c) The temperature profile at 0100 (solid curve), 0700 (long dashed) and 1300 (short dash) hours as calculated from the Ames Mars Global Circulation Model [Haberle et al., this issue].

In addition to imaging clouds, IMP made measurements of the sky brightness at several different wavelengths. From these measurements, the optical depth at each wavelength was derived and found to vary from about 'c = 0.45 to 'c = 0.7 throughout the day with typical midday optical depths of about 'c = 0.5 at 0.67 gin. In general, optical depths started higher in the morning ('c = 0.6) and gradually reduced to the midday value in about 4 hours.

On some mornings, blue opacities were slightly greater than red opacities by about 12% [Smith and Lemmon, this issue]. Nighttime opacities of 'c=0.59 and 'c=0.45, calculated by Thomas et al. [this issue] from observations of Arcturus and Vega, suggest that these opacities change little from day to night.

3. Cloud Physics

In this section we present a description of the physical processes controlling Martian ice clouds. We then discuss the overall organization of the model. We outline the analytic expressions used to describe the processes of sedimentation, nucleation, and condensation. The numerical algorithms used within the simulation are described by Toon et al. [1988]. Similar adaptations of this aerosol model to Mars have been presented by Murphy et al. [ 1990], for simulations of the decay of dust storms, and Michelangeli et al. [1993], for simulations of water clouds. Three fundamental processes dictate the outcome of the simulation: nucleation, condensation, and sedimentation.

The process of nucleation describes the initial formation of an ice crystal spontaneously (homogeneous nucleation) or on a dust grain or similar aerosol (heterogeneous nucleation). Since homogeneous nucleation requires very high supersaturations, in this simulation we consider only the heterogeneous nucleation of ice on dust. Nucleation is highly selective for particle size and depends greatly on the contact parameter m and the supersaturation S = s-l, where the saturation s is the ratio of the partial pressure of water vapor to the vapor pressure of ice. The rate at which nucleation can occur, J, may be expressed as [Pruppacher and Klett, 1978]

4x:2 r•Zs e 2 exp[- AF/kT] J - (2ff In kT) •/2 a• c• where rN is the radius of the dry nucleus, e is the partial pressure of water vapor, c• is the concentration of adsorbing water molecules, T the temperature, and k is Boltzman's constant. Zs, the Zeldovitch factor, is a correction to the equilibrium equation for nucleation that is typically of the order of 0.1. The germ radius ag is given by

ag • 2Mo

9tTp ln(S + 1)

where M is the molecular weight of water, c• is the surface free energy of ice, •R the gas constant,/9 the density of water, and S the supersaturation. The prefactor in front of the exponential is frequently termed the 'kinetic coefficient' with units of a flux and is on order of--- 10 •7 r2v 2 cm '2 s -• for Mars.

AF, the energy of formation onto a flat surface, is

4] •F= •cr [(2+mXl_m)2/ 3

where m is the contact parameter. For a particle of finite size, such as a sphere of radius r•v, the expression for the energy of formation is much more involved, as it requires a more complicated surface geometry [Pruppacher and Klett, 1978, equation. 9-21,22].

COLAPRETE ET AL.: CLOUD FORMATION AT PATHFINDER 9045

An expression that demonstrates the proportionality of the nucleation rate to the supersaturation, the contact parameter, and the particle radius is

J 2 exp[- A(2 + mX1- m)2/[ln(S + 1)12 ] In this expression, A is

A - 16/ro'3M 2/3F2p2kT3

Unfortunately, each of the parameters in A is difficult to measure making the uncertainty in the rate of nucleation very high. Figure 2 shows how the heterogeneous nucleation rate for a spherical surface varies with particle radius, supersaturation, and contact angle. In general, for a given radius and contact angle the nucleation increases with supersaturation. Since no particles smaller than a given size and supersaturation will nucleate, it is ultimately the nucleation which determines the lower end of the size distribution of cloud particles. For higher values of m, nucleation will occur at lower supersaturation [Michela•igeli et al. i993].

The contact parameter m, the ratio of free energy differences between water vapor, ice and dust, can be thought of in a physical sense as the amount of contact between the dust grain and the water cluster. The contact parameter is determined by a mechanical equilibrium such that there is no net force component along the surface of the solid. For a contact parameter of rn = 1, every water molecule which comes into contact with the dust grain will be nucleated. The value of rn varies widely with the type of material nucleating. For example, the contact parameter for silver iodide, a popular cloud seeding agent, is rn = 0.956- 0.988. For liquid water nucleating onto a typical soil, rn = 0.36 - 0.42 [Pruppacher and Klett, 1978]. The roughness and microscopic makeup of the nucleating surface can dramatically change the value of m. Terrestrial clays, a possible analog to the Martian dust, are very susceptible to absorbing impurities that can

increase their nucleation efficiency. For instance, once ice has formed on a dust grain, the dust is said to have been preactivated. With preactivation the contact parameter can increase substantially [Roberts and Hallett, 1968]. Michelangeli et al. [1993] assumed a contact parameter rn -• 0.925 and rn -• 0.975 without and with preactivation respectively for Martian dust. Given our choice of rn = 0.95, Figure 2b shows that dust grains larger than 0.01 to 0.1 gm will be quickly nucleated for modest supersaturations.

Once nucleation occurs the water vapor abundance is quickly reduced by condensation bringing the supersaturation down. As the supersaturation approaches zero the nucleation rate also lowers, effectively stopping further nucleation. Because of its strong dependence on the supersaturation, nucleation ultimately limits the number of ice particles and the size of the ice particles that can form within a cloud.

After nucleation occurs, the process of condensation controls the further growth of ice particles. Condensation is sensitive to the supersaturation, particle size, ability of the particle to be supplied with water vapor, and the exchange of heat with the environment. An approximate growth rate of a particle due to condensation is [Toon et al. 1989]

dt= gonva[(S + l)-4: (l + & g2Qaa)]/[l + & gonoo] In this expression, rtva p is the vapor pressure of water and n•, is the number density of water molecules in the gas phase. Ak is the Kelvin correction that accounts for the effects of particle curvature on the vapor pressure:

-exp(2Mo/pr7).

Qrad, the radiative heating rate of a particle, was assumed to be zero for these calculations.

The terms go, g• and g2 are growth kernels that contain the dependencies of the condensation growth on the water vapor

Log Nucleation rate, J (sec-1) Log Nucleation Rate (J) (sec 0.1

ø). 0 ß 08 II

II

ß 0 0 4

• -?.1 -4.3 •0. 2 0

-2 -1 0 1 2 -2 -1 0 1

Log10 Particle Radius (gm) Log10 Particle Radius (gm)

Figure 2. Plot of the log of the heterogeneous nucleation rate J ( s -•) for a spherical surface showing (a) the dependence on the contact parameter m, and (b) the dependence on the supersaturation S, and the particle radius r•v.

9046 COLAPRETE ET AL.' CLOUD FORMATION AT PATHFINDER

Condensation <• /q[ Evaporation

Cloud (N,r,z) • 50 bins

Sedimentation

Water

Vapor (C,z)

Nucleation

.Evaporation Dust (N,r,z) 50 bins

Sedimentation

Figure 3. Model schematic showing the various processes at work within the model.

diffusion, effects of sedimentation on the flux of mass to the

particle, and the latent heat of condensation'

go - OF, a4/pA

Mpr/SR T gl KF t

g2 - 1/L 47t'r2/9

M is the molecular weight of water, and A is Avogadro's number (r and p have the same meaning as above). Fv and Ft are ventilation corrections that account for the effects of

sedimentation on the flux of water vapor and heat to and from the particles. Le is the latent heat of sublimation, and T is the temperature. In these expressions, D and K are the corrected diffusion and the thermal conduction coefficients, respectively (see Toon et al. [1989] for a more complete description of these terms).

Lastly, the particles settle under the influence of gravity. For low Reynolds numbers the particle fall velocities are calculated from the Stokes relationship:

V/a.- B2pr 2 g/(9rlf )

In this expression, p is the particle density, g is the gravitational acceleration, q is the dynamic viscosity of air, and r•v is the radius of a sphere that has the same volume as the particle. Throughout this work, we assume spherical particles. B is a correction for both particle shape and the high mean free path at low pressures [Toon et al., 1989].

The sedimentation process determines the ability of the particles to transport materials vertically during their lifetime.

Since the fall time is roughly proportional to the inverse of air density and to the inverse of particle radius (since Bocl/pr), high altitudes will be depleted of larger particles in relatively short times [Michelangeli et al., 1993]. As the particles fall into regions of warmer temperatures, the supersaturation can become less than 1, and evaporation will occur.

A fourth process of possible importance is the coagulation of small particles as they collide through BrownJan motion. The coagulation time, or the amount of time required for half of the particles to coagulate, is proportional to the inverse product of the diffusion coefficient, the particle radius, and the number of particles. At the small particle number concentrations encountered in the Martian atmosphere, coagulation does not have a significant effect on the cloud properties.

4. Model Description

The model grid consists of 60 altitude bins each 1.5 km thick, and 50 particle bins with radii ranging from 0.002 to 80 gm for each of the three particle types: dust, water ice, and dust cores within ice particles. Figure 3 shows the general structure of the model and how the physical processes described above lead to interactions among these three particle types.

Three temperature profiles are used throughout this work. The first two (cases A and B) apply a fixed diurnal cycle to the measured Mars Pathfinder profile. The third case (case C) uses temperatures calculated by the Ames Mars Global Circulation Model [Haberle et al., this issue].

Cases A and B assume that the temperature profile measured by ASI/MET represents the coldest nighttime temperatures in the model. While thermal tides weaken this assumption at certain

lOO

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1 10 100 1000

Radius (gm) Radius (gm)

Figure 4. The size distribution of dust after 4 days at three altitudes: z = 5 km, 30 km, and 45 km.

COLAPRETE ET AL.' CLOUD FORMATION AT PATHFINDER 9047

altitudes [Haberle et al., this issue], the purpose of this work is to explore the properties of the clouds as they would form, given the measured profile.

Due to difficulties associated with motions caused by the Pathfinder parachute, the ASI/MET measurements only extended down to about 8 km. From 8 km we interpolate the temperature profile to the ground using subsequent measurements from the Pathfinder lander. Diurnal changes are included for the first two cases as a prescribed offset from this temperature minimum using the relation

T(z,t) - Tmin (Z)-}-a(z)[Sin[2m] +

The offset •z) is defined in a similar fashion to Michelangeli et al. [1993]. The radiative-convective model calculations of Colburn et al. [1989] are used to define a maximum deviation from the nighttime minimum temperature as a function of altitude. At time t = 0, the temperature profile is offset to higher temperatures such that at all altitudes the water vapor supersaturation is less than zero. This offset ensures model stability and allows the dust to come into a steady state without the presence of condensation. The term fl(t) reduces the atmosphere temperature until the profile matches the assumed minimum temperatures measured by Pathfinder. (During this period, there is no diurnal oscillation, i.e., tx(z) = 0.) In our model this occurs over a period of about five Martian days. Once this minimum is reached, the diurnal oscillation is turned on and

the atmospheric temperature oscillates with an amplitude between 2 and 20 K, depending on altitude.

One of the significant differences between the Viking entry profiles and the Pathfinder profile is the inversion measured near 15 kin. To explore the sensitivity of cloud formation at all altitudes to the presence of such an inversion, a second temperature profile was constructed by eliminating this inversion while maintaining the rest of the significant profile features. Figure 1 shows the three temperature profiles used in this study with maximum diurnal variations.

An initial water vapor abundance equal to 1 ! precipitable microns (pt gin) [Jakosky and Farmer, 1982], uniformly mixed

with altitude, is assumed for all three cases. Initially, the dust was assumed to be distributed vertically with a scale height of 12.5 km. The dust was also assumed to have a lognormal size distribution with a modal radius of ro = 0.6 gm. Using an index of refraction for red wavelengths (•,=0.67 gm) of 1.51 + i0.003 [Ockert-Bell et al., 1997], the amount of dust was adjusted so that it had an initial visible optical depth of 'r = 0.5 at •, = 0.67 •m. For blue wavelengths (•, = 0.44 •m) an index of refraction equal to 1.51 + i0.034 was used [Ockert-Bell et al., 1997]. The optical depth was calculated using stratified spherical Mie theory [Ackerman and Toon, 1981; Wiscombe, 1996] at the two visible wavelengths of 0.44 and 0.67 •m. These two wavelengths correspond to filters used by IMP to measure optical depths.

Transport of both water vapor and particles is calculated using an eddy diffusion coefficient that varies as the inverse of the air density such that at the ground q( = 2 x 10 s cm 2 s 'l and at 90 km q( = 6 x 107 cm 2 s -1. No transport is allowed across either the top or bottom boundaries. Both dust and ice particles are allowed to settle into the lowest vertical layer but are not removed from the model entirely. Effectively, this boundary condition assumes the surface is a source of water and dust, offsetting sedimentation. This boundary condition has little effect on our results. After approximately four Martian days the dust distribution and water vapor distributions come into a steady state. In this study we have defined steady state as the point when, bating any cloud formation, the dust particle and water vapor distributions change less than about 3% over the course of a day.

5. Model Results and Discussion

We first (case A) explore the possibility of cloud formation given the unmodified Pathfinder entry temperature profile (Figure 1 a). Before steady state is reached, alterations to the initial dust distribution occur under the influence of sedimentation. After the

fourth day of the model run, further modification of the dust and water vapor distributions occur as condensation is allowed. The largest dust particles fall quickly from high altitudes, leaving only particles smaller than about 0.2 gm above 50 km after 4 days (Figure 4). The loss of large particles has a strong effect on the

Case A

4O

10

-- -- -- Case B

Case C

0.001 0.01 0.1 1 10

Cloud Number Density (cm -3)

Figure 5. Integrated cloud particle concentration as a function of altitude for the three cases simulated in this work. The data shown here are representative of the clouds near their maximum extent in terms of total column number.

9048 COLAPRETE ET AL.' CLOUD FORMATION AT PATHFINDER

subsequent formation of clouds at high altitudes (> 50 km), since the nucleation rate depends strongly on the nucleating particle's radius (Figure 2). For clouds to form by heterogeneous nucleation above about 55 km, dust must be injected into these altitudes at a rate greater than the sedimentation loss rate. Also, since the time it takes for a 1 gm diameter particle at 55 km to fall 10 km is only about 1 hour, clouds forming at these altitudes quickly fall, taking their dust cores with them. This process would deplete the upper atmosphere of dust in just a few days, making further cloud formation here difficult. As ice particles grow they fall quickly and redistribute the water to lower altitudes. The largest increases in water vapor mixing ratios occur just below the locations of cloud formation in areas where ice particles are evaporating. This process also inhibits the formation of clouds above 55 km, since water vapor is trapped at the cloud deck.

Within the constraints of the Mars Pathfinder temperature data, clouds form at two altitudes. Figure 5 shows the location of clouds near their maximum extent (largest ice crystal column concentration). Figure 6 shows the vertical distribution of particles for three altitudes near the maximum extent of the clouds. The higher altitude clouds form within two regions of lower temperatures, corresponding to the temperature inversions at approximately 35 and 45 km (Figure 1). At lower altitudes a cloud forms within the lowest inversion at approximately 8 km. In general, the high altitude clouds have fewer and smaller particles than the lower altitude cloud. This is due primarily to the lack of large condensation nuclei (CN) and the decrease in water vapor concentrations at higher altitudes. In case A the effective particle size, reff, defined as the average particle size weighted by the particle area, for the high altitude clouds is about reff = 1.5 gm at midnight. The lower altitude cloud particles have

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........ 42 km

ß II . .::....: . . .t...% ..........

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Radius ([tm) Radius ([tm)

Figure 6. Cloud particle size distributions at three altitudes at which clouds form for each of the three cases A, B, and C described in the text.

COLAPRETE ET AL.: CLOUD FORMATION AT PATHFINDER 9049

• 3o

,o 20

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1. 1. 1.

1.4 1.4

1.4

6 12

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[8 Time (hrs)

shown in Figure 9. As ice particles grow, the single scatter albedo increases at all wavelengths. This increase in single scatter albedo makes the cloud more neutral, thus making the clouds of larger particles more likely to take on the predominant color of any incident light. However, the cloud/dust particle scattering is never entirely neutral, even at the clouds' maximum extent, but always tends to scatter slightly more red than blue radiation. This effect may explain the light pink color of clouds imaged by IMP. The light blue clouds imaged by IMP may be the result of small CO2 ice particles at high altitude [Clancy and Sandor, 1997]. An alternate explanation may be that the blue clouds are water ice clouds at altitudes that are predominantly free of dust ( > 45 km in this model). The sky color may be blue at such dust-free altitudes. A third possibility is that the particles are very small and thus scatter more efficiently at the shorter blue wavelengths. The merit of each of these possibilities will be discussed within the limits of this study later in the paper.

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q 49 • 1.49 ß 1.49

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Time (hrs)

Figure 7. The effective particle sizes for case A (top) , B (middle), and C (bottom) as a function of altitude and time.

an effective radius of about reff = 2.5 gm at the same time during the night (Figure 7).

The optical depth for case A is shown for two visible wavelengths, ;k = 0.67 gm and ;k = 0.44 gm, in Figure 8. The low-altitude cloud is responsible for 50% of the total optical depth at midnight. The single scatter albedo of the clouds is

1.4

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a20 58

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0 0.54

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---- 0.6 • Case C 0.44 Pm

6 12 18 24

Figure 8. The optical depth at the ground for case A, B, and C at two wavelengths of ;k = 0.67 gm and ;k = 0.44 gm. Local noon is at 12.

9050 COLAPRETE ET AL.: CLOUD FORMATION AT PATHFINDER

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40

30

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06:00

0.6 0.7 0.8 0.9 1 0.6 0.7 0.8 0.9 Single Scatter Albedo Single Scatter Albedo

Figure 9. The single scatter albedo as a function of altitude for case A, B, and C at three times' 1200, 2400 and 0600 hours. The left column is for a wavelength of ;k = 0.67 gm, and the right column is for a wavelength of ;k = 0.44 mm.

The second case (case B) utilizes the modified Pathfinder• temperature profile (Figure lb). The location of the clouds for this case is shown in Figure 5. Given 11 pr gm of water vapor, no clouds form at altitudes below 20 km. Only for water vapor column abundance's greater than about 15 pr gm will a low- altitude cloud form. Even then, this cloud forms at the ground rather than at 8 - 10 km and contains far fewer particles than did the clouds in case A. With the absence of the lowest temperature inversion, water vapor, which was previously trapped at low levels in case A, mixes to higher altitudes. Because of this increase in water vapor the high altitude clouds in case B are slightly thicker physically and have larger particles than the high altitude clouds in case A (Figure 7). The effective cloud particle radius in case B is about ref f = 1.65 gm (Figure 7). The presence

of a low altitude cloud (below 20 km) alters the water distribution

enough to significantly change the particle properties at higher altitudes.

The optical depths for the modified Pathfinder temperature profile (case B) are shown in Figure 8. These optical depths much more closely resemble those calculated from the IMP observations than did the optical depths in case A. The single scatter albedo for case B is shown in Figure 9. As in case A, the formation of cloud particles in case B increases scattering at all visible wavelengths relative to dust. The ratio of blue (X = 0.44 gm) to red (X = 0.67 gm) ground optical depths as a function of time for cases A and B are shown in Figure 10. Case A has a relatively higher blue optical depth than red optical depth when clouds are present. Figure 11 shows the scattering and absorption

COLAPRETE ET AL.' CLOUD FORMATION AT PATHFINDER 9051

1.03

1.02

1.01

1

0.99

0.98

0.97

0.96

...... Case A t "• I '"! I !

Case B

I • I i/xl ' I I I I I I I I

t .... i .... • , , , , i , , , , i ....

0 5 10 15 20

Time (hrs)

Figure 10. The ratio of blue 0• = 0.44 gm) to red 0• = 0.67 gm) ground optical depths as a function of time for cases A and B. Case C is very similar to case B.

efficiencies of cloud particles for both case A and B and for both red and blue wavelengths. Also shown in Figure 11 is the volume fraction of ice on each cloud particle. For case A the peak absorption efficiency occurs for particles at about 2 gm and is higher for the blue than for the red wavelengths. However, the scattering efficiency of particles at this size is nearly equal for both red and blue wavelengths. Thus the increase in blue absorption efficiency results in a greater increase in the total blue extinction efficiency, relative to red. In case B there is also an increase in the absorption efficiency; however, since the particles are smaller, the scattering efficiencies are not equal. Decreases in blue scattering efficiencies offsetting increases in blue absorption efficiencies lead to little net change in the total extinction efficiency. Figure 12 shows the total extinction efficiencies for red and blue wavelengths for cases A and B. Also plotted is the particle size distribution (dN/d log (r)) for the same altitude and time. The larger particle sizes of clouds simulated in case A lead to greater blue opacities relative to red opacities.

Lastly, we explore the properties of clouds that form given the third temperature profile (case C). A cloud forms at a similar altitude for case C as for cases A and B (Figure 5); however, its time and location are dictated strongly by the thermal tides. As in case B, no clouds form below approximately 20 km and no clouds form above 50 kin; however, clouds are present throughout the day. Clouds forming during the night occur at about 40 kin, while the thickest clouds forming during the day occur at about 20 kin. As the thermal tides propagate vertically, clouds form at varying locations under a wider range of supersaturations and with a wider range of dust particle sizes in case C. This variation leads to a broad size distribution, as is seen

in Figure 6. In general, the clouds are composed of fewer particles than those in cases A and B.

The continued production of cloud particles throughout the day leads to a nearly constant effective particle size with time (Figure 7) that closely matches those calculated by Smith and Letomort [this issue]. However, the optical depth (Figure 8), while agreeing initially in the morning hours, increases toward noon in a way that was not observed. This increase is due to a cooling at about 25 km due to the thermal tide. Also seen in Figure 8 is that the blue optical depth is never greater than the red optical depth. This is due to the continued presence of cloud particles throughout the day and their broad size distribution. Neglected in this simulation is the effect of solar warming of cloud particles. It is likely that as the elevation of the Sun increases, warming of cloud particles would lead to sublimation and a decrease in cloud particle numbers. This effect might reduce the noontime optical depth and match the observed data more closely. It is the goal of our current work to include these radiative effects along with the microphysical model.

The single scatter albedo for case C is shown in Figure 9. As in the other two cases, the clouds lead to an increase in the single scatter albedo, making the clouds a more neutral scatterer than the surrounding dust. However, for case C the increase in single scatter albedo is not as great as for cases A or B. This is primarily due to the fewer number of particles in the case C cloud compared to cases A and B.

Case B

,. • 10 x Shell

5 5 l/Volume Fraction • 4 ]P• '•'•Q Red 10 x •ell • • 4 , • Q, Blue/ • '' Volume Fraction • • / •S' Red/ • •

•2 2

0.01 0.1 I 10 0.0• 0.• • •0

Pa•icle Radius (•) Pmticle Radius (•)

Figure 11. The scattering efficiency Qs and the absorption efficiency Qa for the lower altitude cloud in case A (particles at midnight and an altitude of 9 kin) and the higher altitude cloud in case B (particles at midnight and an altitude of 35 kin). The water shell volume fraction is also plotted for both cases A and B. Results for case C are very similar to those of case B.

9052 COLAPRETE ET AL.: CLOUD FORMATION AT PATHFINDER

4.5

4

3.5

3

12.5 2

1.5 1

0.5

0

0.01

Case A

/ /

// :! Particle Radius (Ixrn)

4.5

3.5

3

2.,5 2

1.5 1

0.5

0

O.Ol

Case B

/ / i/ :',

0.1 1 10

Patic]e Radius (!b-n)

Figure 12. The total extinction efficiency for the lower altitude cloud in case A (particles at midnight and an altitude of 9 km) and cloud in case B (particles at midnight and an altitude of 35 km). Also plotted is the particle size distribution (dN/d log (r)) for both cases A and B. Results for case C are very similar to those of case B.

6. Conclusions

With the Pathfinder temperature profile (case A), the model described here produces clouds at two altitudes. If the inversion at 8 km measured by Pathfinder is included in the model, clouds form at 10 km and 35 km. If the inversion is removed (case B), the clouds form only at altitudes above 35 km, but are more vertically extensive. The presence of the lower inversion affects the particle size distribution by creating larger particles at both high and low altitudes. Using the diurnal temperature profile calculated by the Ames Mars general circulation model (MGCM) a primary cloud formed between 20 and 40 km and was present throughout the day. IMP observed early morning optical depths of •: = 0.57 at all wavelengths, which is more consistent with case B (no temperature inversion) and case C. However, case B slightly overestimates the nighttime optical depth, and case C overestimates the noontime optical depth. Solar warming of cloud particles and a resulting sublimation of the particles may reduce the cloud number at around noontime. This effect could

bring the case C results into close agreement with the observations made by IMP.

On some days, IMP observed a measurable difference in the optical depth between blue and red wavelengths, and a slightly higher optical depth (•: = 0.67). Our case A not only has greater optical depth, but it also has blue opacities, which are slightly higher than red opacities. However, our case A cloud has too large an optical depth to be consistent with the IMP data. Neither cases B nor C simulated a higher blue-to-red opacity. In these cases the peak in the size distribution did not correspond to higher blue extinction efficiencies due to the smaller particle sizes associated with the clouds. Possibly, the ratio of blue to red opacity is affected by factors other than cloud altitude such as water vapor abundance or nucleation rates. We will consider these possibilities in a future paper.

The color of the clouds as observed by IMP is a product of the relative intensities of light either passing through or scattering off the cloud particles. If the particles are neutrally scattering, then they will reflect all wavelengths nearly equally. If directly illuminated by the Sun, these clouds appear white. The background sky color can provide contrast so that the cloud color

stands out more clearly. Such is the case on Earth, as clouds reflect solar illumination against a background dominated by Rayleigh scattering. For light passing through a cloud, the relative amount of light scattered or absorbed at each wavelength determines the color of the cloud. While the composition of

small particles plays some role, the dominate factor in determining their total extinction is their size and shape. A good example of this is the blue Sun phenomenon seen here on Earth. Particles with an average size of about 1-1.5 gm will create a relative increase in red-to-blue absorption, causing sunlight passing through to appear bluish. If, however, the particles are slightly smaller than this, with an average size of about 0.5 I.tm, the Sun will appear red as more blue light is scattered away from the viewer [Horvath et al., 1994]. In either case, relatively narrow size distributions are required as a broad distribution of particles will contribute all across the extinction curve and effectively wash out any color discrimination.

With this in mind, we can review the three scenarios described earlier with regard to the blue and pink clouds observed by IMP. Just as described above, small particles directly illuminated by the Sun could scatter more blue than red to an observer and create the effect of a blue cloud. This is the case for CO2 clouds

forming at very high altitudes (>75 km) as proposed by Clancy and Sandor [1998]. For such clouds to appear blue, not only do they have to scatter more blue than red initially, the blue light that is scattered must also make its way to the observer through the Martian atmosphere. While the nature of CO2 nucleation is not well known, if it were to occur heterogeneously, nucleating particles, such as dust, would have to be present in sufficient quantities and sizes. It is not clear that this is the case for altitudes above 75 km.

If the blue clouds consist of water ice particles, a difficulty remains in that, while still relatively small, they likely do not occur at altitudes high enough to be directly illuminated in the predawn hours. If these clouds are occurring at altitudes predicted by this work, they are being illuminated by diffuse sky light. One possibility is that the high-altitude clouds (-40 km) are just at the top of the dust layer and are illuminated by Rayleigh scattering from above. At these altitudes the predawn Rayleigh scattering optical depth is about x - 0.03 in the blue.

COLAPKETE ET AL.: CLOUD FORMATION AT PATHFINDER 9053

While this opacity is small, it may be large enough for scattered sky light to preferentially illuminate these high-altitude clouds while the Sun is below the horizon, making them appear bluish.

If the cloud particle size distribution is particularly narrow, then these clouds will scatter, or absorb, more of one color

compared to another. This effect is seen in small amounts in the simulations done here (Figure 12) and may also explain the higher blue opacities on some days. A narrow particle distribution may be the result of evaporation or a short sudden burst of nucleation; however, neither of these were recreated in this simulation.

Acknowledgments. This work was supported by NASA grant NAG5-6900. The authors are grateful to Robert Haberle for the use of the Ames Mars GCM temperature data.

References

Ackerman, T. P. and O. B. Toon. Absorption of visible radiation in atmosphere containing mixtures of absorbing and non-absorbing particles, Appl. Opt., 20, 3661-3668, 1981.

Clancy, R. T. and B. J. Sandor, CO2 ice clouds in the upper atmosphere of Mars, Geophys. Res. Lett., 25(4), 489-492, 1998..

Clancy, R. T., A. W. Grossman, M. J. Wolff, P. B. James, D. J. Rudy, Y. N. Billawala, B. J. Sandor, S. W. Lee and D. O. Muhleman, Water vapor saturation at low altitudes around Mars aphelion: A key to Mars climate?, Icarus, 122, 36-62, 1996.

Colburn, D. S., J. B. Pollack and R. M. Habede, Diurnal variations in the optical depth at Mars, Icarus, 79, 159-189, 1989.

Haberle, R. M., et al., General circulation model simulation of the Mars Pathfinder atmospheric structure instrument / meteorology data, J. Geophys. Res., this issue.

Horvath, H., G. Metzig, O. Preining and R. F. Pueschel, Observation of a blue sun over New Mexico, U.S.A., on 19 April 1991, Atmos. Environ. 28(4), 621-630, 1994.

Jakosky, B. M., The role of seasonal reservoirs in the Mars water cycle. II. Coupled models of the regolith, the polar caps, and atmospheric transport, Icarus, 55, 19-39, 1983.

Jakosky, B. M., and C. B. Farmer, The seasonal and global behavior of water vapor in the Mars atmosphere: Complete global results of the Viking atmospheric water detector experiment, J. Geophys. Res., 87, 2999-3019, 1982.

Jakosky, B. M., R. W. Zurek, and M. R. LaPointe, The observed day-to- day variability of Mars atmospheric water vapor, Icarus, 73, 80-90, 1988.

James, P. B., The role of water-ice clouds in the M•.rtian hydrologic cycle, J. Geophys. Res., 95, 1439-1445, 1990.

Kahn, R., Ice haze, snow, and the Mars water cycle, J. Geophys. Res., 95, 14,677-14,693, 1990.

Magalhaes, J. A., J. T. Schofield, and A. Seiff, Results of the Mars Pathfinder Atmospheric Structure Investigation, J. Geophys. Res., this issue.

Michelangeli, D. V., O. B. Toon, R. M. Haberle, and J. B. Pollack, Numerical simulation of the formation and evolutii•n of water ice

clouds in the Martian atmosphere, Icarus, 100, 261-285, 1993. Murphy, J. R., O. B. Toon, R. M. Haberle, and J. B. Pollack, Numerical

simulation of the decay of Martian global dust storms, J. Geophys. Res., 95, 14,629-14,648, 1990.

Ockert-Bell, M. E., J. F. Bell III, J.P. Pollack, C. P. McKay, and F. Forget, Absorption and scattering properties of the Martian dust in the solar wavelengths, J. Geophys. Res., 102, 9039-9050, 1997.

Pollack, J. B., C. B. Leovy, P. W. Greiman, and Y. Mintz, A Martian general circulation experiment with large topography, J. Atmos. Sci., 38, 3-29, 1981.

Pruppacher, H. R., and J. D. Klett, Microphysics of Clouds and Precipitation, D. Reidel, Norwell, Mass., 1978.

Robets, P., and J. Hallett, A laboratory study of ice nucleation properties of some material particulates, Q. J. R. Meteorol. Soc., 94, 25-34, 1968.

Smith, P. H., and M. Lemmon, Opacity of the Martian atmosphere measured by the Imager for Mars Pathfinder, J. Geophys. Res., this issue.

Thomas, N., D. T. Britt, K. E. Herkenhoff, S. L. Murchie, B. Semenov, H. U. Keller, and P.H. Smith, Observations of Phobos, Deimos, and bright stars with the Imager for Mars Pathfinder, J. Geophys. Res., this issue (a).

Toon, O. B., R. P. Turco, D. Wesphall, R. Malone, and M. S. Liu, A multidimensional model for aerosols: Description of computational analogs, J. Atmos. Sci., 45, 2123-2143, 1988.

Toon, O. B., R. P. Turco, J. Jordan, J. Goodman, and G. Ferry, Physical processes in polar stratospheric ice clouds, J. Geophys. Res., 94, 11,359-11,380, 1989.

Wiscombe, W. J., Mie scattering: advances in technique and fast, vector speed computer codes, NCAR/TN-140+STR, Natl. Cent. For Atmos. Res., Boulder Colo., 1996.

A. Colaprete, Laboratory for'Atmospheric and Space Physics, University of Colorado, Boulder CO, 80309 (e-mail:tony@sunra.colorado.edu).

O. B. Toon, Laboratory for Atmospheric and Space Physics, University of Colorado, Boulder CO, 80309.

J. A. Magalhfies, Space Sciences Division, NASA Ames Research Center, Moffett Field, CA, 94035.

(Received May 11, 1998; revised August 24, 1998; accepted September 30, 1998.)