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The Martian upper atmosphere structure from the Viking spacecraft experiments
ICAaUS36, 189--197 (1978)
The Martian Upper Atmosphere Structure from the Viking Spacecraft Experiments M. N. I Z A K O V Space Research Institute, Academy of Sciences, Moscow, USSR Received August 26, 1977; revised April 17, 1978 Height profiles of the atmospheric component concentrations measured by the onboard mass spectrometers during the Viking lander descents can be well described under the assumption of a smoother temperature profile and smaller eddy diffusion coefficients at altitudes of 120 to 170 km than those determined by McElroy et al. The eddy diffusion coefficient in the Martian thermosphere during the Viking experiments was equal to 5 X 107 cm~sec-1, according to our interpretation. 1
During the descent of the Viking 1 and 2 landers in the M a r t i a n a t m o s p h e r e excellent mass spectrometer m e a s u r e m e n t s were carried o u t ; the d a t a published b y Nier et al. (1976), Nier and M c E l r o y (1976, 1977), and M c E l r o y et al. (1976), together with those from the m e a s u r e m e n t s in the lower a t m o s p h e r e (Nier et al., 1976; Seiff and Kirk, 1976) contribute m u c h to our knowledge a b o u t the structure of the M a r t i a n atmosphere. T h e M a r t i a n a t m o spheric composition and processes were studied using this d a t a (Yung et al., 1977; K o n g and M c E l r o y 1977). However, in our opinion some details of the d a t a processing given b y M c E l r o y et al. (1976) require additional analysis, and some conclusions are objectionable.
T h e height profile of t e m p e r a t u r e , T(h), calculated b y M c E l r o y et al. (1976) a n d Nier and M c E l r o y (1977) from t h a t of the C02 concentration, nco~(h), measured with a mass s p e c t r o m e t e r has the sinuous shape that, in the opinion of M c E l r o y et al.,
reflects the real variations of t e m p e r a t u r e associated with the propagation of tidal waves in the atmosphere. This m a y be p a r t l y so; however it also m a y be due p a r t l y to errors in the concentration m e a surements and t e m p e r a t u r e calculations. T h e error in t e m p e r a t u r e obtained f r o m the concentration profile is ~T ~ + 6 0 ° at 200 km, according to the estimates of M c E l r o y et al.; it then rapidly decreases during the descent and at 140 to 160 k m it is only ~T ~ 4-5 to 10°; then it increases again up to ~T ~ 20 to 30 ° at 120 km. To obtain such a high accuracy of t e m p e r a t u r e m e a s u r e m e n t s at 140 to 160 km, the concentrations should be measured with an accuracy of 5n~ ~ 3 to 5%. Is this possible in flight ? I t has been assumed b y M c E l r o y et al. (1976) t h a t 6n~ can be determined from the difference in the concentration values for one c o m p o n e n t calculated f r o m various mass peaks (for CO2 t h e y are 44, 22, 12). Here it is implied t h a t the following errors can be neglected as c o m p a r e d with those of calibration: errors due to the inflight instability of e q u i p m e n t p a r a m e t e r s (ion source, especially); due to the calcu189 0019-1035/78/0362-0189502.00/0 Copyright O 1978 by Academic Press, Inc. All rights of reproduction in any form reserved.
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M.N. IZAKOV
lation of concentrations in the atmosphere from those in the system (the noncontrolled variations in the angle of a t t a c k can contribute to this error); and, for the components such as N2 and CO, falling within the same peaks, due to the additional calculations required for their separation. But if all these errors are really negligible, such a strong and nonmonotonic dependence of 8T on altitude cannot then be explained. The additional analysis of these questions, t h a t is within the competence of the experimenters only, should be of great importance. Usually in-flight concentration measurement errors are about 15-30°~o (Engebretson et al., 1977; K i r b y Docken and Oppenheimer, 1977). When temperature is calculated from concentration with the barometric formula, as in M c E l r o y et al. (1976), it is determined, in principle, by the logarithmic derivative of concentration; and numerical differentiation of the empirical curve plotted with an error m a y decrease the accuracy of the results. F r o m this viewpoint it is expedient to integrate numerically the concentration curve in order to obtain pressure, as has been done b y Seiff and Kirk (1976), and then to obtain temperature from it. If the sufficiently accurate calculations b y this method are used, the relative error of temperature obtained can remain the same as t h a t of concentration. The interpolation between experimental points--since t h e y are far from each o t h e r - is required both for differentiation and integration of the n,(h) profile. If the error of concentration measurements 3n, is known, it is expedient, instead of interpolating, to plot a smoothed approximating curve (e.g., b y the method of least squares) deviating from experimental points b y not more than the error value. This m a y help exclude artificial waves from the curve n.(h). T h e nco2(h) profiles calculated with constant temperature versus altitude differ from experimental points b y not more than
30% (Fig. 1), excluding lower points at h < 130 k m which could be effected b y pressure saturation in the mass spectrometer (McElroy et al., 1976), and the upper one at h = 199 k m from Viking 2. So we conclude that during the Viking 1 experiment the Martian thermosphere temperature was T~ = 180 4- 40°K, and during the Viking 2, To = 130 4- 30°K. If the structure of the n,(h) and T(h) profiles really reflects the influence of atmospheric waves, the smoothed curve should be close to the equilibrium profile with no waves. This profile is the most i m p o r t a n t feature of the atmosphere, the first to be determined. 3
The profile of an eddy diffusion coefficient K(h) was defined b y M c E l r o y et al. (1976) and Nier and M c E l r o y (1977) from the profiles of argon and nitrogen concentrations, nAt(h) and nx2(h), and temperature T(h), calculated using the profile neon(h). The values obtained, K = (1 to 4) X 109 cm 2 sec-1 at 170 kin, were v e r y large (Fig. 2). F r o m our viewpoint these values of K were calculated with too large an error and at 140-170 km t h e y were overestimated for the following reasons. In M c E l r o y et al. (1976), K was found from the equation 0 Inn, --. On
0 In T +
-
-
Oh
D , / H . ° q- K / I 4 -
,
(1)
D, + K
t h a t is, the barometric formula in a differential form for a transition layer between the homosphere and the heterosphere. Here, 0 In n,/On + 0 In T/On = - - 1 / H , ; H,~ is the scale height of the a-th component pressure ; H , ° -= R o T / M,g is the scale height of the ~-th component pressure for diffusive equilibrium in the heterosphere; and /4 = RoT/~Ig is the mean scale height for diffusive equilibrium, t J Different authors interpret differently the term "scale height." It is reasonable to define it as an
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I t is obvious t h a t in the homosphere where K >> D, and in the heterosphere where D~ >> K, Eq. (1) transforms into the common barometric formula with /7 and H , °, respectively. E q u a t i o n (1) is derived from the continuity equation for a long-lived atmospheric component (i.e., the one for which the characteristic time of chemical reactions is much longer than t h a t of diffusion) if the latter is assumed to have a stationary distribution and thermal diffusion is neglected (see, e.g., Izakov, 1976). However, if acoustic-gravitational waves propagate in the atmosphere, the distributions are not stationary and it is then impossible to neglect the time derivative On,/Ot in the continuity equation; and, probably, it is also impossible to neglect the thermodiffusive term for large temperature gradients; and, consequently, (1) is inapplicable. Let us assume t h a t the measured profiles of n,(h), T(h) are close to stationary or that the smoothing procedure can be used so t h a t t h e y are made to approach the stationary distributions. Then, and only then, (1) is applicable. However, even in this case, K is derived from (1) with a v e r y high error. Actually, if (1) is solved for K it can be written as
D ~ ( 1 / H . ° -- 1 / H , ) K
=
1/H~ -- 1/171
(2)
I t is clear that K is a fraction, the numerator and denominator of which are the differences of the values with significant errors [-since they include a logarithmic derivative of the empirical profile n.(h), temperatures calculated from it and its logarithmic derivatives~ ; besides scale heights in (2) do not differ much in their a l t i t u d e i n t e r v a l over w h i c h a c e r t a i n v a l u e c h a n g e s e times, i.e., 1 / H o = [(1/4~)04~/0h]. I t is o b v i o u s
t h a t scale h e i g h t s for p r e s s u r e a n d c o n c e n t r a t i o n are different: 1 / H p = 1 / H n -- 0 In T / O h . H o w e v e r , t h e t e r m "scale h e i g h t " is o f t e n u s e d to describe t h a t for p r e s s u r e in t h e case of diffusive e q u i l i b r i u m Hp ° = RoT/Mg.
EXPERIMENTS
191 TABLE I
PARAMETERS FOR THE CALCULATION OF BINARY DIFFUSION COEFFICIENTS OF ATMOSPHERIC GASES INTO CO~ FROM THE EQUATION D ~ = A , z T s , ~ exp ( -- B , ~ / T ) n -1
a-~
A.o ( X 10-17)
S.~
B.~
CO-CO2 O~-CO2 N2-CO~ Ar-CO2 H2-CO~ He-CO2
0. 424 1.15 2.28 1.28 2.30 2.45
0. 803 0.661 0.570 0. 646 0. 750 0. 720
-61.3 113.6 89.1 11.7 --
values since the state is close to equilibrium and the zone considered is not far from the homopause (where K = De b y definition). I t is fully apparent t h a t the resulting error in determining K can be v e r y large. But there are other factors: besides the error due to the structure of the expression for K and the errors in the parameters it includes, another error is added due to inaccuracies in the value of the average molecular mass 21~ = ~ M . n . / ~ n ~ in /t. In fact, within the height range discussed, atomic oxygen becomes one of the basic components of the Martian atmosphere, the concentration of which cannot be measured reliably b y a mass spectrometer but is determined (with some error) from models. According to the models, the concentration of no is compared with t h a t of nco~ at 180 to 200 k m (Liu and Donahue, 1976; Izakov and Krasickij, 1976; Kong and McElroy, 1977). I t can be assumed t h a t due to the above mentioned errors the values of D. at high altitudes " d r a g " with them the values of K in the calculations presented b y McE l r o y et al. (1976) and Nier and M c E l r o y (1977); indeed, it is seen from Fig. 2 t h a t the dependence K(h) at altitudes higher than 130 k m is v e r y similar to the well known dependence of D,(h).
192
M. N. IZAKOV
h(Km) 200
i
h(Km) 200
160
120 106
IO’
108
FIG. 1. Height profilw of concentrations of Viking 1 data, (b) from the Viking 2 data. o--cOz, A--Ar, q-Nz, @--oz. 11esults cm2/sec-1; -.-.-, profiles 2’(h), K(h) taken K(h) taken from RIcElroy et al. (1976).
109
10’0
N.
(cm-‘)
the Martian atmospheric components. (a) from the Experimental points (Nier and hlcb%wy, 1976) : of calculations: p, constant T, K = 5 X 107 from hlcElroy et al. (19i6); -- --, constant Y,
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And there is one more important point. It is seen from (1) that if the molecular mass of some component M, is equal to the average molecular mass i@ and, t’hereforc, Ha0 = A, the equation degenerates and it is impossible to derive K from it. But if M, is close to Al, it is also practically impossible to determine K since very large differences in K result in very small differences in n,(h). This is the case for argon since MA? = 40 is close to a = 43.5. Hence, it is not expedient to determine K from the curve n*,(h); it is better to use this curve for deriving T(h), though an error in the homosphere will be somewhat larger than from ma,(h) due to a greater difference of molecular weight from the a,verage. In summary we can say that to determine rrliably the eddy diffusion coefficient K by the method discussed, first it is necessary that the profile of the concentration of a component, with molecular mass strongly differing from the average should be used (the profile for helium would bc ideal) and, second, reliable measurements of atomic oxygen should be simultaneously available. 4
The following calculation was made to check the considerations discussed above. According t’o Nier and McElroy (1976), the concentrations of components, n,, were taken at some initial height ho near the lower boundary of the measurement domain and n,(h) was calculated from the barometric formula written in the form n, (h)= n,
xexP(
T(ho)
(ho)T(h) j-1
&(&+;)dh’).
(3)
In this case the molecular diffusion coefficients D((Y, COz) were calculated according to the data in Mason and Marrcro (1970), where a great number of laboratory
193
EXPERIMENTS
measurements dependence
was
generalized
and
the
Da0 = AeoTSaficxp ( - B,o/ T) n-l was given, where Ao,a,Sap, Bap are constants for the given pair of molecules (see Table I). The eddy diffusion coefficient K was chosen within the limits determined by the theoretical estimations (Izakov, 1977) and then it was made more accurate during the calculations. Several versions were calculated: (a) with constant T (T = 180°K for Viking 1 and T = 130°K for Viking 2 results and K = 5 X 10’ cm2 set--’ ; (b) with the profiles T(h) and K(h) taken from McElroy et al. (1976); and (c) with constant T and K(h) from McElroy et aZ. (1976) [the auxiliary version t’o find out how K effects n,(h) (see Fig. l)]. Comparison of versions (a) and (b) shows that profiles obtained with constant T and K bett’er describe the experimental data. K values can rapidly decrease at altitudes higher than 13(r170 km; i.e., there can be a turbopause as on the Earth. Here D, >> K always, so K values do not in practice effect n,(h) profiles. Comparison of versions (b) and (c) shows how little the n,(h) profiles differ for a large difference in K values, for argon especially, as follows from the consideration discussed above. 5
Now let us see how the data from the Viking spacecraft compares with that from other experiments and models. According to the above interpretation of the Viking spacecraft measurements the most probable temperature of the Martian upper thcrmosphcrc is T, = 130 =t 30°K in the morning (about 09 hours local time) and T, = 180 f 40°K in the daytime (about 16 hours local time). If the experimental data from which T, can be determined (using scale heights for hydrogen, carbon dioxide, and
194
M. N. IZAKOV ,
h(Km)
,
i
• I
150 -
,
,
~ J
-1 "
®
J ' ~
-
I ..~/
IO0
/
./
/ !
50
~ O
~
i05
,'1"i I0 6
*
,
;
i°, ° ,
!
,
I0 7
,
,' ....
I I0 °
,
,
,' ....
I I0 9
i
,
,
I
i ,ll
I0 I°
K, ONa-coa (cm2/s) FIG. 2. Height profiles of the eddy diffusion coefficient K(h) and the molecular diffusion coefficient D~2_co2(h) in the M a r t i a n atmosphere, if) K values from Viking 1 results according to McElroy et al. (1976) ; Ax K values from the Viking 2 results according to McElroy et al. (1976). .... , K value describing well the Viking spacecraft data according to the present paper; . . . . . , limits of possible variations of K in the M a r t i a n atmosphere according to Izakov (1977) and Golitsyn (1973); . . . . , dependence K(h) recommended by McElroy and Kong (1976); X - X - X and . . . . . . . are the dependence of the molecular diffusion coefficients D(N2_CO2) for Vikings 1 and 2, respectively.
plasma) are summarized, the dependence of T . on the solar activity is obvious, as we h a v e already shown (Izakov, 1973, 1976). T h e global mean t e m p e r a t u r e of the u p p e r t h e r m o s p h e r e is within T . = 350 to 400°K for the high solar activity corresponding to the solar radio flux at 10.7 cm, Fi0.7 = (150 to 200) X 10-22 W m -2 Hz-1; and T~ = 270 to 280°K for the m o d e r a t e l y low solar a c t i v i t y corresponding to F10.7 = 100 X 10 - ~ W m -2 Hz -1. True, sometimes this dependence is obviously screened b y additional sources of heating in the thermosphere t h a t can be caused b y the dynamical effect of the lower a t m o s p h e r e via acoustic-gravitational waves propa-
gating upwards and dissipating in the thermosphere, on the one hand, and b y the energy transfer f r o m the solar wind through the magnetosphere, on the other. I n a n y case the theoretical model ( I z a k o v et al., 1976), in which the main heat source due to solar ultraviolet radiation, the infrared heat sink, and heat removal b y molecular heat conductivity were t a k e n into account, gives T . close to the empirical values (Izakov, 1973). Therefore, the uv-solar absorption is really the main source of heat in the low-latitude t h e r m o s p h e r e during the quiet periods. During the Viking 1 and 2 experiments there was a fairly low level of solar activity,
VIKING SPACECRAFT EXPEI~IMENTS corresponding to F10.7 ~ 70 X 10 -22 W m -e H z -1 (Solar Geophys. Data, 1976). If this fact is t a k e n into account as well as the fact t h a t Mars was near its aphelion (which leads to a solar flux decrease b y a b o u t 14% compared to when M a r s is at its m e a n distance f r o m the Sun), it is found according to the theoretical model ( I z a k o v et al., 1976) t h a t T~ should be 200 to 220°K at this period. We can see t h a t this value is nevertheless higher t h a n T~ f r o m the Viking spacecraft data. H o w can we explain an even cooler t h e r m o s p h e r e ? We can do it b y taking into account t u r b u l e n t heat flux downwards from the lower thermosphere. T u r b u lence in the lower thermosphere produces a heat source due to dissipation of t u r b u l e n t energy on the one hand, and a heat sink due to t u r b u l e n t heat conductivity on the other (Hunten, 1974; Johnson, 1975; Izakov, 1978). These effects are p a r t l y compensated ; however, under certain conditions one of t h e m can prevail. I t depends on the value of the critical flux Richardson n u m ber, and it was argued t h a t turbulence cools the M a r t i a n thermosphere when the troposphere is quiet a n d heats it when the troposphere is disturbed (Izakov, 1978). As has been described above, the fast increase of K with height from (2 to 4) X 107 cm 2 sec - I at 100 k m up to (1 to 4) X 109 cm 2 sec -1 at 170 k m (Fig. 2) was obtained from the Viking spacecraft d a t a b y M c E l r o y et al. (1976). Alternately, the same d a t a are well described according to our interpretation with a K at 100 to 150 k m t h a t does not change with height, equal t o K = 5 X 107cm ~sec -I. Let us compare these conclusions with the other data on the eddy diffusion coefficient in the M a r t i a n u p p e r atmosphere. N e a r the t u r b o p a u s e on M a r s K can v a r y within (0.9 to 8) X 107 cm 2 sec -~ according to our estimate, in the context of the similarity t h e o r y using the R i c h a r d s o n O b u k h o v law (Izakov, 1977). E s t i m a t e s b y a slightly different m e t h o d b y Golitsyn a n d
195
Steklov (1977) give K = 8 × 106 cm 2 sec -1, close to our lower limit. Using our estimates (Izakov, 1977) and Golitsyn's (1973) t h a t resulted in K = (0.1 + 4) X 106 cm ~ sec -1 near the M a r t i a n surface and interpolating between t h e m (assuming the linear dependence of log K on h), we plotted the profiles K(h) shown in Fig. 2. According to our d a t a the h o m o p a u s e is at 120 to 140 km. A model of the M a r t i a n atmospheric composition has been calculated based on these profiles ( I z a k o v and Krasickij, 1977), the various versions of which a d e q u a t e l y describe the experimentally measured variations of concentrations for some components. On the other hand, the model of Liu and D o n a h u e (1976) where a heightindependent K = 4 X l0 s cm 2 sec -~ was used, leads to values of the O and CO concentrations near the ionospheric maxim u m significantly lower t h a n those obtained experimentally. I n M c E l r o y and K o n g (1977) the profile K(h) was given without a n y discussion (it is also shown in Fig. 2), with the values of K m u c h larger t h a n ours. Our experience in the composition model calculation allows us to conclude t h a t these K values are also overestimated. Indeed, the profiles K(h) r e c o m m e n d e d are v e r y a p p r o x i m a t e ; so, for example, there is p r o b a b l y a t u r b o p a u s e at some height over the h o m o p a u s e (at 150 to 170 km), as on the Earth, a b o v e which K decreases rapidly, since on M a r s as on the E a r t h turbulence in the stable lower thermosphere is maintained b y the energy of acoustical-gravitational waves coming from below and dissipating in the lower t h e r m o s p h e r e (Hines, 1965; Hodges, 1967; Gavrilov and Shved, 1975). However, it will not in practice effect the profiles of concentrations of components since here D~ >> K under all conditions. 6
T h e following conclusions can thus be made.
196
M. N. IZAKOV
T h e w a v e s on t h e h e i g h t profiles of c o n c e n t r a t i o n s of a t m o s p h e r i c c o m p o n e n t s a n d a t m o s p h e r i c t e m p e r a t u r e can b e d u e p a r t l y to a c o u s t i c a l - g r a v i t a t i o n a l waves, p a r t l y to t h e errors in m e a s u r e m e n t s a n d c a l c u l a t i o n s . T o s e p a r a t e real a n d artificial w a v e s , a d e t a i l e d a n a l y s i s of errors is r e q u i r e d . T h e t e m p e r a t u r e profiles d u r i n g the Viking experiments could have been smoother than those determined by McE l r o y et al. (1976). T h e d e t e r m i n a t i o n of t h e e d d y diffusion coefficient u s i n g t h e c o n c e n t r a t i o n profiles is p o s s i b l e o n l y if c e r t a i n c o n d i t i o n s a r e m e t : s t a t i o n a r y d i s t r i b u t i o n s of c o n c e n t r a t i o n s ; t h e m o l e c u l a r m a s s of t h e c o m p o n e n t u s e d sufficiently d i f f e r e n t f r o m t h e a v e r a g e m o l e c u l a r m a s s ( h e l i u m is t h e b e s t choice) ; s u f f i c i e n t l y a c c u r a t e d e t e r m i n a t i o n of c o n c e n t r a t i o n s of all t h e c o m p o n e n t s , i n c l u d i n g a t o m i c o x y g e n . T h e v a l u e s of t h e t u r b u l e n t diffusion coefficient K o b t a i n e d b y M c E l r o y et al. (1976) a t 100 to 120 k m a r e close to r e a l b u t a t h i g h e r a l t i t u d e s they are overestimated, the error increasing w i t h h e i g h t . T h e v a l u e s of K d u r i n g t h e V i k i n g e x p e r i m e n t a t 100 to 150 k m were e q u a l to 5 X 107 c m 2 sec -1 a n d a t h i g h e r a l t i t u d e s t h e y a r c l i k e l y to decrease. L o w t e m p e r a t u r e s in t h e t h e r m o s p h e r e d u r i n g t h e V i k i n g s p a c e c r a f t mission (130 4 - 3 0 ° K a t 09 h o u r s local t i m e "rod 180 4-40°K a t 16 h o u r s local t i m e ) a r e e x p l a i n e d a b e v e all b y t h e low level of s o l a r a c t i v i t y a n d , to a s m a l l e r e x t e n t , b y M a r s b e i n g n e a r its a p h e l i o n a n d , p r o b a b l y to a c e r t a i n degree, b y t u r b u l e n t h e a t flux from the thermosphere. M e a s u r e m e n t s of t h e h e i g h t profiles of the atmospheric component concentrations, sufficiently c o m p l e t e a n d a c c u r a t e , a r e one of t h e b e s t m e t h o d s of i n v e s t i g a t i n g t h e s t r u c t u r e a n d d y n a m i c s of p l a n e t a r y atmospheres. REFERENCES ENGEBRETSON~ M. Z., MAUERSBERGER~ S., AND
POTTER W. E.
(1977). Extension of atomic
nitrogen measurements in the lower thermosphere. J. Geophys. Res. 82, 3291-3294. GAVRILOV,N. M., ANDSHVED, G. M. (1975). On the closure of the equation system for the turbulized layer of the upper atmosphere. Ann. Geophys. 31, 375-388. GOLITSYN, G. S. (1973). Vvedenije v dinamiku planetnyh atrnosfer (Introduction to planetary atmosphere dynamics). Leningrad, Gidrometeoizdat. GOLITSYN, G. S., AND STEKLOV, A. F. (1977). Elementary consideration of vertical atmospheric structure. Doklady A N SSSR 232, 786-789. Hines, C. O. (1965). Dynamical heating of the upper atmosphere. J. Geophys. Res. 70, 177 183. HODGES, R. R., JR. (1967). Generation of turbulence in the upper atmosphere by internal gravity waves, J. Geophys. Res. 72, 3455~3458. Hunten, D. M. (1974). Energetics of thermospheric eddy transport, J. Geophys. Res. 79, 2533 2534. IZAKOV, M. N., (1973). On the Martian thermosphere temperature. Kosmich. l~sled. 11, 761-766. IZAKOV, M. N. (1976). Structure and dynamics of the upper atmospheres of Venus and Mars. Uspehi Fiz. Nauk 119, 295-342 (Soy. Phys. Usp. 19, 503-529). IZAKOV, M. N. (1977). Evaluation of eddy mixing coefficient and of homopause altitude on Venus, Mars and Jupiter. Kosmich. lssled. 15, 248-254. IZAKOV, M. N. (1978). Turbulence effect on the thermal regime of planetary thermospheres. Kosmich. I~sled. 16, 403-411. IZAKOV, M. N., KRASICKIJ, O. E., AND MOROZOV, S. K. (1976). Dependence of structure of Mars' and Venus' thermospheres on solar activity. Kosmich. IssIed. 14, 585 591. IZAKOV, M. N., AND KnASICK~J,O. P. (1977). Model of Martian atmosphere composition. Kosmich. Issled. 15, 455-469. JOttNSON~ F. S. (1975). Transport processes in the upper atmosphere. J. Atmos. SeN. 32, 1658-1662. KmBY--DocKEN, K., AND OPPENHEIMER, M. (1977). Indirect determination of molecular oxygen densities in the daytime thermosphere from atmosphere. Explorer C composition measurements. J. Geophys. Res. 82, 3503 3508. KONG, T. Y., AND McELRoY, M. B. (1977). Photochemistry of the Martian atmosphere. Icar~ls 32, 168-189. LIu, S. C., AND DONAHUE, T. M. (1976). The regulation of hydrogen and oxygen escape from Mars. Icarus 28, 231-246. MASON, E. A., AND MAanERO, T. R. (1970). The diffusion of atoms and molecules. Adv. Atom. ~}lolec. Phys. 6, 15& 167. McELRoY, M. B., KONG, T. Y., YUNC., Y. L., AND NIER, A. O. (1976). Composition and structure
VIKING SPACECRAFT E X P E R I M E N T S of the Martian upper atmosphere: Analysis of results from Viking. Science 194, 1295 1298. MCELROY, M. B., AND KONG, T. Y. (1976). Oxidation of the Martian surface: Constraints due to chemical processes in the atmosphere. Geophys. Res. Lett. 3, 569-572. NIER, A. O., HANSON, W. B., SEIFF, A., McELRoY, M. B., SPENCER, N. W., DUCKETT,R. J., KNIGHT, T. C. D., AND COOK, W. S. (1976). Composition and structure of the Martian atmosphere: Preliminary results from Viking 1. Science 193, 786-788. NIER, A, O., AND McELRoY, M. B. (1976). Structure of the neutral upper atmosphere of Mars: Results
197
from Viking 1 and Viking 2. Science 194, 12981300. Nml~, A. O., AND McELRo~, M. B. (1977). Composition and structure of Mars' upper atmosphere : Results from the neutral mass spectrometers on Viking 1 and 2. J. Geophys. Res. 82, 4341-4350. SEIFF, A., AND KIRK, D. B. (1976). Structure of Mars' atmosphere up to 100 kilometers from the entry measurements of Viking 2. Science 194, 1300--1303. Solar Geophysical Data (1976). N 38% Pt. 1. YUNG, Y. L., STROBEL~D. F., KONG~ T. Y.j AND MCELRO¥, M. B. (1977). Photochemistry of nitrogen in the Martian atmosphere. Icarus 30, 26-41.
The Martian Upper Atmosphere Structure from the Viking Spacecraft Experiments M. N. I Z A K O V Space Research Institute, Academy of Sciences, Moscow, USSR Received August 26, 1977; revised April 17, 1978 Height profiles of the atmospheric component concentrations measured by the onboard mass spectrometers during the Viking lander descents can be well described under the assumption of a smoother temperature profile and smaller eddy diffusion coefficients at altitudes of 120 to 170 km than those determined by McElroy et al. The eddy diffusion coefficient in the Martian thermosphere during the Viking experiments was equal to 5 X 107 cm~sec-1, according to our interpretation. 1
During the descent of the Viking 1 and 2 landers in the M a r t i a n a t m o s p h e r e excellent mass spectrometer m e a s u r e m e n t s were carried o u t ; the d a t a published b y Nier et al. (1976), Nier and M c E l r o y (1976, 1977), and M c E l r o y et al. (1976), together with those from the m e a s u r e m e n t s in the lower a t m o s p h e r e (Nier et al., 1976; Seiff and Kirk, 1976) contribute m u c h to our knowledge a b o u t the structure of the M a r t i a n atmosphere. T h e M a r t i a n a t m o spheric composition and processes were studied using this d a t a (Yung et al., 1977; K o n g and M c E l r o y 1977). However, in our opinion some details of the d a t a processing given b y M c E l r o y et al. (1976) require additional analysis, and some conclusions are objectionable.
T h e height profile of t e m p e r a t u r e , T(h), calculated b y M c E l r o y et al. (1976) a n d Nier and M c E l r o y (1977) from t h a t of the C02 concentration, nco~(h), measured with a mass s p e c t r o m e t e r has the sinuous shape that, in the opinion of M c E l r o y et al.,
reflects the real variations of t e m p e r a t u r e associated with the propagation of tidal waves in the atmosphere. This m a y be p a r t l y so; however it also m a y be due p a r t l y to errors in the concentration m e a surements and t e m p e r a t u r e calculations. T h e error in t e m p e r a t u r e obtained f r o m the concentration profile is ~T ~ + 6 0 ° at 200 km, according to the estimates of M c E l r o y et al.; it then rapidly decreases during the descent and at 140 to 160 k m it is only ~T ~ 4-5 to 10°; then it increases again up to ~T ~ 20 to 30 ° at 120 km. To obtain such a high accuracy of t e m p e r a t u r e m e a s u r e m e n t s at 140 to 160 km, the concentrations should be measured with an accuracy of 5n~ ~ 3 to 5%. Is this possible in flight ? I t has been assumed b y M c E l r o y et al. (1976) t h a t 6n~ can be determined from the difference in the concentration values for one c o m p o n e n t calculated f r o m various mass peaks (for CO2 t h e y are 44, 22, 12). Here it is implied t h a t the following errors can be neglected as c o m p a r e d with those of calibration: errors due to the inflight instability of e q u i p m e n t p a r a m e t e r s (ion source, especially); due to the calcu189 0019-1035/78/0362-0189502.00/0 Copyright O 1978 by Academic Press, Inc. All rights of reproduction in any form reserved.
190
M.N. IZAKOV
lation of concentrations in the atmosphere from those in the system (the noncontrolled variations in the angle of a t t a c k can contribute to this error); and, for the components such as N2 and CO, falling within the same peaks, due to the additional calculations required for their separation. But if all these errors are really negligible, such a strong and nonmonotonic dependence of 8T on altitude cannot then be explained. The additional analysis of these questions, t h a t is within the competence of the experimenters only, should be of great importance. Usually in-flight concentration measurement errors are about 15-30°~o (Engebretson et al., 1977; K i r b y Docken and Oppenheimer, 1977). When temperature is calculated from concentration with the barometric formula, as in M c E l r o y et al. (1976), it is determined, in principle, by the logarithmic derivative of concentration; and numerical differentiation of the empirical curve plotted with an error m a y decrease the accuracy of the results. F r o m this viewpoint it is expedient to integrate numerically the concentration curve in order to obtain pressure, as has been done b y Seiff and Kirk (1976), and then to obtain temperature from it. If the sufficiently accurate calculations b y this method are used, the relative error of temperature obtained can remain the same as t h a t of concentration. The interpolation between experimental points--since t h e y are far from each o t h e r - is required both for differentiation and integration of the n,(h) profile. If the error of concentration measurements 3n, is known, it is expedient, instead of interpolating, to plot a smoothed approximating curve (e.g., b y the method of least squares) deviating from experimental points b y not more than the error value. This m a y help exclude artificial waves from the curve n.(h). T h e nco2(h) profiles calculated with constant temperature versus altitude differ from experimental points b y not more than
30% (Fig. 1), excluding lower points at h < 130 k m which could be effected b y pressure saturation in the mass spectrometer (McElroy et al., 1976), and the upper one at h = 199 k m from Viking 2. So we conclude that during the Viking 1 experiment the Martian thermosphere temperature was T~ = 180 4- 40°K, and during the Viking 2, To = 130 4- 30°K. If the structure of the n,(h) and T(h) profiles really reflects the influence of atmospheric waves, the smoothed curve should be close to the equilibrium profile with no waves. This profile is the most i m p o r t a n t feature of the atmosphere, the first to be determined. 3
The profile of an eddy diffusion coefficient K(h) was defined b y M c E l r o y et al. (1976) and Nier and M c E l r o y (1977) from the profiles of argon and nitrogen concentrations, nAt(h) and nx2(h), and temperature T(h), calculated using the profile neon(h). The values obtained, K = (1 to 4) X 109 cm 2 sec-1 at 170 kin, were v e r y large (Fig. 2). F r o m our viewpoint these values of K were calculated with too large an error and at 140-170 km t h e y were overestimated for the following reasons. In M c E l r o y et al. (1976), K was found from the equation 0 Inn, --. On
0 In T +
-
-
Oh
D , / H . ° q- K / I 4 -
,
(1)
D, + K
t h a t is, the barometric formula in a differential form for a transition layer between the homosphere and the heterosphere. Here, 0 In n,/On + 0 In T/On = - - 1 / H , ; H,~ is the scale height of the a-th component pressure ; H , ° -= R o T / M,g is the scale height of the ~-th component pressure for diffusive equilibrium in the heterosphere; and /4 = RoT/~Ig is the mean scale height for diffusive equilibrium, t J Different authors interpret differently the term "scale height." It is reasonable to define it as an
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I t is obvious t h a t in the homosphere where K >> D, and in the heterosphere where D~ >> K, Eq. (1) transforms into the common barometric formula with /7 and H , °, respectively. E q u a t i o n (1) is derived from the continuity equation for a long-lived atmospheric component (i.e., the one for which the characteristic time of chemical reactions is much longer than t h a t of diffusion) if the latter is assumed to have a stationary distribution and thermal diffusion is neglected (see, e.g., Izakov, 1976). However, if acoustic-gravitational waves propagate in the atmosphere, the distributions are not stationary and it is then impossible to neglect the time derivative On,/Ot in the continuity equation; and, probably, it is also impossible to neglect the thermodiffusive term for large temperature gradients; and, consequently, (1) is inapplicable. Let us assume t h a t the measured profiles of n,(h), T(h) are close to stationary or that the smoothing procedure can be used so t h a t t h e y are made to approach the stationary distributions. Then, and only then, (1) is applicable. However, even in this case, K is derived from (1) with a v e r y high error. Actually, if (1) is solved for K it can be written as
D ~ ( 1 / H . ° -- 1 / H , ) K
=
1/H~ -- 1/171
(2)
I t is clear that K is a fraction, the numerator and denominator of which are the differences of the values with significant errors [-since they include a logarithmic derivative of the empirical profile n.(h), temperatures calculated from it and its logarithmic derivatives~ ; besides scale heights in (2) do not differ much in their a l t i t u d e i n t e r v a l over w h i c h a c e r t a i n v a l u e c h a n g e s e times, i.e., 1 / H o = [(1/4~)04~/0h]. I t is o b v i o u s
t h a t scale h e i g h t s for p r e s s u r e a n d c o n c e n t r a t i o n are different: 1 / H p = 1 / H n -- 0 In T / O h . H o w e v e r , t h e t e r m "scale h e i g h t " is o f t e n u s e d to describe t h a t for p r e s s u r e in t h e case of diffusive e q u i l i b r i u m Hp ° = RoT/Mg.
EXPERIMENTS
191 TABLE I
PARAMETERS FOR THE CALCULATION OF BINARY DIFFUSION COEFFICIENTS OF ATMOSPHERIC GASES INTO CO~ FROM THE EQUATION D ~ = A , z T s , ~ exp ( -- B , ~ / T ) n -1
a-~
A.o ( X 10-17)
S.~
B.~
CO-CO2 O~-CO2 N2-CO~ Ar-CO2 H2-CO~ He-CO2
0. 424 1.15 2.28 1.28 2.30 2.45
0. 803 0.661 0.570 0. 646 0. 750 0. 720
-61.3 113.6 89.1 11.7 --
values since the state is close to equilibrium and the zone considered is not far from the homopause (where K = De b y definition). I t is fully apparent t h a t the resulting error in determining K can be v e r y large. But there are other factors: besides the error due to the structure of the expression for K and the errors in the parameters it includes, another error is added due to inaccuracies in the value of the average molecular mass 21~ = ~ M . n . / ~ n ~ in /t. In fact, within the height range discussed, atomic oxygen becomes one of the basic components of the Martian atmosphere, the concentration of which cannot be measured reliably b y a mass spectrometer but is determined (with some error) from models. According to the models, the concentration of no is compared with t h a t of nco~ at 180 to 200 k m (Liu and Donahue, 1976; Izakov and Krasickij, 1976; Kong and McElroy, 1977). I t can be assumed t h a t due to the above mentioned errors the values of D. at high altitudes " d r a g " with them the values of K in the calculations presented b y McE l r o y et al. (1976) and Nier and M c E l r o y (1977); indeed, it is seen from Fig. 2 t h a t the dependence K(h) at altitudes higher than 130 k m is v e r y similar to the well known dependence of D,(h).
192
M. N. IZAKOV
h(Km) 200
i
h(Km) 200
160
120 106
IO’
108
FIG. 1. Height profilw of concentrations of Viking 1 data, (b) from the Viking 2 data. o--cOz, A--Ar, q-Nz, @--oz. 11esults cm2/sec-1; -.-.-, profiles 2’(h), K(h) taken K(h) taken from RIcElroy et al. (1976).
109
10’0
N.
(cm-‘)
the Martian atmospheric components. (a) from the Experimental points (Nier and hlcb%wy, 1976) : of calculations: p, constant T, K = 5 X 107 from hlcElroy et al. (19i6); -- --, constant Y,
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And there is one more important point. It is seen from (1) that if the molecular mass of some component M, is equal to the average molecular mass i@ and, t’hereforc, Ha0 = A, the equation degenerates and it is impossible to derive K from it. But if M, is close to Al, it is also practically impossible to determine K since very large differences in K result in very small differences in n,(h). This is the case for argon since MA? = 40 is close to a = 43.5. Hence, it is not expedient to determine K from the curve n*,(h); it is better to use this curve for deriving T(h), though an error in the homosphere will be somewhat larger than from ma,(h) due to a greater difference of molecular weight from the a,verage. In summary we can say that to determine rrliably the eddy diffusion coefficient K by the method discussed, first it is necessary that the profile of the concentration of a component, with molecular mass strongly differing from the average should be used (the profile for helium would bc ideal) and, second, reliable measurements of atomic oxygen should be simultaneously available. 4
The following calculation was made to check the considerations discussed above. According t’o Nier and McElroy (1976), the concentrations of components, n,, were taken at some initial height ho near the lower boundary of the measurement domain and n,(h) was calculated from the barometric formula written in the form n, (h)= n,
xexP(
T(ho)
(ho)T(h) j-1
&(&+;)dh’).
(3)
In this case the molecular diffusion coefficients D((Y, COz) were calculated according to the data in Mason and Marrcro (1970), where a great number of laboratory
193
EXPERIMENTS
measurements dependence
was
generalized
and
the
Da0 = AeoTSaficxp ( - B,o/ T) n-l was given, where Ao,a,Sap, Bap are constants for the given pair of molecules (see Table I). The eddy diffusion coefficient K was chosen within the limits determined by the theoretical estimations (Izakov, 1977) and then it was made more accurate during the calculations. Several versions were calculated: (a) with constant T (T = 180°K for Viking 1 and T = 130°K for Viking 2 results and K = 5 X 10’ cm2 set--’ ; (b) with the profiles T(h) and K(h) taken from McElroy et al. (1976); and (c) with constant T and K(h) from McElroy et aZ. (1976) [the auxiliary version t’o find out how K effects n,(h) (see Fig. l)]. Comparison of versions (a) and (b) shows that profiles obtained with constant T and K bett’er describe the experimental data. K values can rapidly decrease at altitudes higher than 13(r170 km; i.e., there can be a turbopause as on the Earth. Here D, >> K always, so K values do not in practice effect n,(h) profiles. Comparison of versions (b) and (c) shows how little the n,(h) profiles differ for a large difference in K values, for argon especially, as follows from the consideration discussed above. 5
Now let us see how the data from the Viking spacecraft compares with that from other experiments and models. According to the above interpretation of the Viking spacecraft measurements the most probable temperature of the Martian upper thcrmosphcrc is T, = 130 =t 30°K in the morning (about 09 hours local time) and T, = 180 f 40°K in the daytime (about 16 hours local time). If the experimental data from which T, can be determined (using scale heights for hydrogen, carbon dioxide, and
194
M. N. IZAKOV ,
h(Km)
,
i
• I
150 -
,
,
~ J
-1 "
®
J ' ~
-
I ..~/
IO0
/
./
/ !
50
~ O
~
i05
,'1"i I0 6
*
,
;
i°, ° ,
!
,
I0 7
,
,' ....
I I0 °
,
,
,' ....
I I0 9
i
,
,
I
i ,ll
I0 I°
K, ONa-coa (cm2/s) FIG. 2. Height profiles of the eddy diffusion coefficient K(h) and the molecular diffusion coefficient D~2_co2(h) in the M a r t i a n atmosphere, if) K values from Viking 1 results according to McElroy et al. (1976) ; Ax K values from the Viking 2 results according to McElroy et al. (1976). .... , K value describing well the Viking spacecraft data according to the present paper; . . . . . , limits of possible variations of K in the M a r t i a n atmosphere according to Izakov (1977) and Golitsyn (1973); . . . . , dependence K(h) recommended by McElroy and Kong (1976); X - X - X and . . . . . . . are the dependence of the molecular diffusion coefficients D(N2_CO2) for Vikings 1 and 2, respectively.
plasma) are summarized, the dependence of T . on the solar activity is obvious, as we h a v e already shown (Izakov, 1973, 1976). T h e global mean t e m p e r a t u r e of the u p p e r t h e r m o s p h e r e is within T . = 350 to 400°K for the high solar activity corresponding to the solar radio flux at 10.7 cm, Fi0.7 = (150 to 200) X 10-22 W m -2 Hz-1; and T~ = 270 to 280°K for the m o d e r a t e l y low solar a c t i v i t y corresponding to F10.7 = 100 X 10 - ~ W m -2 Hz -1. True, sometimes this dependence is obviously screened b y additional sources of heating in the thermosphere t h a t can be caused b y the dynamical effect of the lower a t m o s p h e r e via acoustic-gravitational waves propa-
gating upwards and dissipating in the thermosphere, on the one hand, and b y the energy transfer f r o m the solar wind through the magnetosphere, on the other. I n a n y case the theoretical model ( I z a k o v et al., 1976), in which the main heat source due to solar ultraviolet radiation, the infrared heat sink, and heat removal b y molecular heat conductivity were t a k e n into account, gives T . close to the empirical values (Izakov, 1973). Therefore, the uv-solar absorption is really the main source of heat in the low-latitude t h e r m o s p h e r e during the quiet periods. During the Viking 1 and 2 experiments there was a fairly low level of solar activity,
VIKING SPACECRAFT EXPEI~IMENTS corresponding to F10.7 ~ 70 X 10 -22 W m -e H z -1 (Solar Geophys. Data, 1976). If this fact is t a k e n into account as well as the fact t h a t Mars was near its aphelion (which leads to a solar flux decrease b y a b o u t 14% compared to when M a r s is at its m e a n distance f r o m the Sun), it is found according to the theoretical model ( I z a k o v et al., 1976) t h a t T~ should be 200 to 220°K at this period. We can see t h a t this value is nevertheless higher t h a n T~ f r o m the Viking spacecraft data. H o w can we explain an even cooler t h e r m o s p h e r e ? We can do it b y taking into account t u r b u l e n t heat flux downwards from the lower thermosphere. T u r b u lence in the lower thermosphere produces a heat source due to dissipation of t u r b u l e n t energy on the one hand, and a heat sink due to t u r b u l e n t heat conductivity on the other (Hunten, 1974; Johnson, 1975; Izakov, 1978). These effects are p a r t l y compensated ; however, under certain conditions one of t h e m can prevail. I t depends on the value of the critical flux Richardson n u m ber, and it was argued t h a t turbulence cools the M a r t i a n thermosphere when the troposphere is quiet a n d heats it when the troposphere is disturbed (Izakov, 1978). As has been described above, the fast increase of K with height from (2 to 4) X 107 cm 2 sec - I at 100 k m up to (1 to 4) X 109 cm 2 sec -1 at 170 k m (Fig. 2) was obtained from the Viking spacecraft d a t a b y M c E l r o y et al. (1976). Alternately, the same d a t a are well described according to our interpretation with a K at 100 to 150 k m t h a t does not change with height, equal t o K = 5 X 107cm ~sec -I. Let us compare these conclusions with the other data on the eddy diffusion coefficient in the M a r t i a n u p p e r atmosphere. N e a r the t u r b o p a u s e on M a r s K can v a r y within (0.9 to 8) X 107 cm 2 sec -~ according to our estimate, in the context of the similarity t h e o r y using the R i c h a r d s o n O b u k h o v law (Izakov, 1977). E s t i m a t e s b y a slightly different m e t h o d b y Golitsyn a n d
195
Steklov (1977) give K = 8 × 106 cm 2 sec -1, close to our lower limit. Using our estimates (Izakov, 1977) and Golitsyn's (1973) t h a t resulted in K = (0.1 + 4) X 106 cm ~ sec -1 near the M a r t i a n surface and interpolating between t h e m (assuming the linear dependence of log K on h), we plotted the profiles K(h) shown in Fig. 2. According to our d a t a the h o m o p a u s e is at 120 to 140 km. A model of the M a r t i a n atmospheric composition has been calculated based on these profiles ( I z a k o v and Krasickij, 1977), the various versions of which a d e q u a t e l y describe the experimentally measured variations of concentrations for some components. On the other hand, the model of Liu and D o n a h u e (1976) where a heightindependent K = 4 X l0 s cm 2 sec -~ was used, leads to values of the O and CO concentrations near the ionospheric maxim u m significantly lower t h a n those obtained experimentally. I n M c E l r o y and K o n g (1977) the profile K(h) was given without a n y discussion (it is also shown in Fig. 2), with the values of K m u c h larger t h a n ours. Our experience in the composition model calculation allows us to conclude t h a t these K values are also overestimated. Indeed, the profiles K(h) r e c o m m e n d e d are v e r y a p p r o x i m a t e ; so, for example, there is p r o b a b l y a t u r b o p a u s e at some height over the h o m o p a u s e (at 150 to 170 km), as on the Earth, a b o v e which K decreases rapidly, since on M a r s as on the E a r t h turbulence in the stable lower thermosphere is maintained b y the energy of acoustical-gravitational waves coming from below and dissipating in the lower t h e r m o s p h e r e (Hines, 1965; Hodges, 1967; Gavrilov and Shved, 1975). However, it will not in practice effect the profiles of concentrations of components since here D~ >> K under all conditions. 6
T h e following conclusions can thus be made.
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M. N. IZAKOV
T h e w a v e s on t h e h e i g h t profiles of c o n c e n t r a t i o n s of a t m o s p h e r i c c o m p o n e n t s a n d a t m o s p h e r i c t e m p e r a t u r e can b e d u e p a r t l y to a c o u s t i c a l - g r a v i t a t i o n a l waves, p a r t l y to t h e errors in m e a s u r e m e n t s a n d c a l c u l a t i o n s . T o s e p a r a t e real a n d artificial w a v e s , a d e t a i l e d a n a l y s i s of errors is r e q u i r e d . T h e t e m p e r a t u r e profiles d u r i n g the Viking experiments could have been smoother than those determined by McE l r o y et al. (1976). T h e d e t e r m i n a t i o n of t h e e d d y diffusion coefficient u s i n g t h e c o n c e n t r a t i o n profiles is p o s s i b l e o n l y if c e r t a i n c o n d i t i o n s a r e m e t : s t a t i o n a r y d i s t r i b u t i o n s of c o n c e n t r a t i o n s ; t h e m o l e c u l a r m a s s of t h e c o m p o n e n t u s e d sufficiently d i f f e r e n t f r o m t h e a v e r a g e m o l e c u l a r m a s s ( h e l i u m is t h e b e s t choice) ; s u f f i c i e n t l y a c c u r a t e d e t e r m i n a t i o n of c o n c e n t r a t i o n s of all t h e c o m p o n e n t s , i n c l u d i n g a t o m i c o x y g e n . T h e v a l u e s of t h e t u r b u l e n t diffusion coefficient K o b t a i n e d b y M c E l r o y et al. (1976) a t 100 to 120 k m a r e close to r e a l b u t a t h i g h e r a l t i t u d e s they are overestimated, the error increasing w i t h h e i g h t . T h e v a l u e s of K d u r i n g t h e V i k i n g e x p e r i m e n t a t 100 to 150 k m were e q u a l to 5 X 107 c m 2 sec -1 a n d a t h i g h e r a l t i t u d e s t h e y a r c l i k e l y to decrease. L o w t e m p e r a t u r e s in t h e t h e r m o s p h e r e d u r i n g t h e V i k i n g s p a c e c r a f t mission (130 4 - 3 0 ° K a t 09 h o u r s local t i m e "rod 180 4-40°K a t 16 h o u r s local t i m e ) a r e e x p l a i n e d a b e v e all b y t h e low level of s o l a r a c t i v i t y a n d , to a s m a l l e r e x t e n t , b y M a r s b e i n g n e a r its a p h e l i o n a n d , p r o b a b l y to a c e r t a i n degree, b y t u r b u l e n t h e a t flux from the thermosphere. M e a s u r e m e n t s of t h e h e i g h t profiles of the atmospheric component concentrations, sufficiently c o m p l e t e a n d a c c u r a t e , a r e one of t h e b e s t m e t h o d s of i n v e s t i g a t i n g t h e s t r u c t u r e a n d d y n a m i c s of p l a n e t a r y atmospheres. REFERENCES ENGEBRETSON~ M. Z., MAUERSBERGER~ S., AND
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VIKING SPACECRAFT E X P E R I M E N T S of the Martian upper atmosphere: Analysis of results from Viking. Science 194, 1295 1298. MCELROY, M. B., AND KONG, T. Y. (1976). Oxidation of the Martian surface: Constraints due to chemical processes in the atmosphere. Geophys. Res. Lett. 3, 569-572. NIER, A. O., HANSON, W. B., SEIFF, A., McELRoY, M. B., SPENCER, N. W., DUCKETT,R. J., KNIGHT, T. C. D., AND COOK, W. S. (1976). Composition and structure of the Martian atmosphere: Preliminary results from Viking 1. Science 193, 786-788. NIER, A, O., AND McELRoY, M. B. (1976). Structure of the neutral upper atmosphere of Mars: Results
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from Viking 1 and Viking 2. Science 194, 12981300. Nml~, A. O., AND McELRo~, M. B. (1977). Composition and structure of Mars' upper atmosphere : Results from the neutral mass spectrometers on Viking 1 and 2. J. Geophys. Res. 82, 4341-4350. SEIFF, A., AND KIRK, D. B. (1976). Structure of Mars' atmosphere up to 100 kilometers from the entry measurements of Viking 2. Science 194, 1300--1303. Solar Geophysical Data (1976). N 38% Pt. 1. YUNG, Y. L., STROBEL~D. F., KONG~ T. Y.j AND MCELRO¥, M. B. (1977). Photochemistry of nitrogen in the Martian atmosphere. Icarus 30, 26-41.