Tile DTIVERSIT'Y J' MICHIGAN
RADIO ZASTRONCOi`Y OBSERVATORY
Report No. 66-3
A MODEL MARTIAN ATMOSPHERE AND IONOSPHERE
Newbern Smith and Abigail Beutler
Sponsored by National Aeronautics and Space Administration
Grant No. NSG 572
March 1966
Department of Astronomy
Department of Electrical Engineerirg

Newbern Smith a bd Abigail E. Beutler; Radio Astronomy Observatory
Th i Uniter-ity of Michigan, Ann Arbor
A Model Martian Atmosphere and Ionosphere
Summary
The results of the Mariner IV experiment indicate that CO2 is probably
the only significant constituent of the lower Martian atmosphere. Assuming
this, we have derived en equilibrium model atmosphere and ionosphere of Pars
for the subsolar region and at sunspot maximum. A temperature profile resulted
from consideration o; radiation balance between the troposphere and mesosphere,
vibrational relaxation of CO2 and the heat released by photodissociation and
ionization processes, Dissociation begins at about 70 km producing maximum
densities of the dissociation products between 80 and 90 km. Owing to the
low density of the atmosphere, CO2 is distributed in diffusive rather than
chemical equilibrium. The photochemistry of the ionosphere includes a multiplicity of charge transfer reactions. The resulting ionosphere is bimodal,
the lower peak being predominantly 02. The upper peak, predominantly 0 +
is due to the transition between chemical and diffusive equilibrium and may
thus be considered an F2 peak. The difference between this ionosphere and
that observed by Mariner IV can be resolved by considering thst the F2 peak
is lacking at high latitude due in part to an expected decrease in temperature
and in part because 0+ decreases much more rapidly than 02 with decreasing
ionizing flux density.

I. Introduction
Thus far no satisfactory model of the Martian atmosphere and ionosphere
has been advanced. Early models (i.e. prior to 1964) (Kuiper 1952;
Grandjean and Goody 1955; Hess 1958; Kellogg and Sagan 1961; Chamberlain 1962)
all. included abundant N2 to account for a considerably overestimated surface
pressure. It appears from the results of Mariner IV that retention of N2
as a significant constituent (e.g. McElroy et al 1965; Chamberlain and
McElroy 1966) can be justified mainly by tradition. Several models have
recognized the importance of chemical equilibrium in the photodissociation
region but have not treated concurrently the diffusion processes. The
resulting models are either purely in chemical equilibrium (e,g. Chamberlain
and McElroy 1966) or purely in diffussive equilibrium (e.g. Gross et al 1965)
neither of which is satisfactory. Many models to not give adequate consideration to heat balance as a determinant of temperature distribution in the
atmosphere; in the resulting models the radiation efficiency is not sufficient
to dispose of the heat input to the atmosphere (e.g. Johnson 1965). In
addition to the obvious atmospheric processes mainly below the mesopause,
an adequate treatment of the atmosphere must take into consideration the
following: photodissociation; chemical equilibrium; diffusive equilibrium;
photoionization; photochemistry of the ionosphere; loss of heat by radiation
and conduction; and the resulting temperature profile, which in turn influences each of the preceeding processes. In addition, it is necessary
for the equation of continuity to be satisfied for the atmosphere as a
whole. In this paper we have tried to consider the simultaneous effect of
these phenomena.
II. The Lower Martian Atmosphere
Kaplan, et al (1964), confirmed by Owen (1964), obtained a CO2 abundance
of 50 + 20 m atm, and their analysis led to a surface composition that can
be described empirically by
1/2
P Pi 19
where pi is the partial pressure of CO2 in mb and p is the total pressure
in mb including a few mb of A and N2. The minimum surface pressure
would then be 7 mb, for a pure CO2 atmosphere.

2
Observations made from Mariner IV during the flyby of July 15, 1965,
(Kliore, et al, 1965) were interpreted to meal a surface scale height of
8.5 km, and a surface pressure of 4 - 6 mb for pure CO2 or 5 - 7 mb of an
equal mixture of CO2 and A. The small scale height would seem to rule out
any significant amount of N2, or else imply a surface temperature radically
lower than for an airless rotating planet of Mar's albedc at its distance
from the sun.
Considering probably uncertainties of observations, it appears that the
lower Martian atmosphere may be entirely composed of CO2 with only traces of
such other gases as A or N2.
The mean daytime surface temperature of Mars is about 240~K and the
temperature variation from subsolar regions to the terminator has been
estimated as about 800K, while the subsolar temperature has been measured
as high as 300~K. A mean subsolar temperature of 280~K and a diurnal range
of 1600K are consistent with the observed mean surface temperature. The
surface temperature at a solar zenith angle of 70~ would then be about
2300K and the air temperature about 50TK less (Mintz 1961) corresponding
to the Mariner IV observation.
We shall assume a surface atmospheric temperature of 2300K and a surface
17 -3(l)
pressure of 7 mb, leading to a surface [CO ] equal to 2.3 x 10 cm
-1
For a pure CO2 atmosphere, the adiabatic rate is -5.2 deg km. Elementary
radiation balance consideration put the tropopause temperature at 1670K, and
consequently its height at about 12 km, with a pressure of 2 mb. Following
the reasoning of Goody (1957) the temperature is assumed to fall off to
about 1530 well above the tropopause,
Pressure and density calculations on this model put the level of vibrational relaxation for C02 at about 60 km, where [C02] = 4 x 101 cm-3. On
the assumption that radiative heat transport is negligible above this region,
and that no heat source is present, an adiabatic lapse rate is once again
assumed, up to the level where [CO2] 1014 cm 3 (about 70 km) and its
dissociation begins.
(1) Throughout this paper, square brackets are used to indicate number
densities,

3
Some of the energy absorbed above this level is converted into heat, all
of which can be disposed of by radiation of CO2 between 70 and 95 km. Nevertheless the dissociation of 02 and CO2 above this level constitutes a heat
source which affects the temperature profile.
III. Photodissociation Region
The photodissociative solar flux at any level depends upon the optical
thickness of the atmosphere through which it has passed. By considering the
dissociation cross section, a, of CO2 and its diffusion coefficient we
concluded that CO2 is always in diffusion. The optical thickness, TCO, at
any height above 95 km is therefore given by
CO (a n H)CO
where
n = number density
and H = local scale height calculated at a mean gravity of 342 cm sec
at 140 km.
A first approximation to the abundances of the photodissociation
products can be calculated by solving the photochemical equilibrium equations
using CO2 in diffusion as the only absorbing constituent. The absorption by
the resulting 02 must now be considered in this iterative procedure.
The dissociative solar flux for CO2 at the distance of Mars is 6 x 1011
-2 -1 -7
photons cm sec at sunspot maximum with a dissociation rate of 1,2 x 10
sec1 (Nawrocki and Papa 1961). The dissociation contiuum begins at 1700 A
-20 2 -18 2
with a cross section of about i0 cm increasing to 10 cm at 1350 A,
-120 2 -17 2
reaching a minimum of 4 x 1020 cm at 120Q A and increasing to 1017 cm
0 12
below 1150 A. The dissociative flux for 02 is about 2 x 101 photons
-2 -1 -6 1
cm sec, with a dissociation rate of 2.5 x 10 sec. The weak Herzberg
continuum does not contribute appreciably to the total dissociation, except
as noted below, compared to the Schumann-Runge bands. The latter begin at
oot1 0 -20ll 2.raigt -17 2
about 1800 A with a cross section of 1020 cm increasing to 10 cm at
0 -.
1400 A and then decreasing to 5 x 10 19 cm below 1250 A. The absorption
by 2 a 0 e 0bion
by CO and 0 in the Hcrzbcg region between 1800 A and 2400 A is very small.

4
This radiation will therefore penetrate to lower levels where it is effective
in dissociating molecular oxygen. The atomic oxygen thus liberated is
available for the oxidation of carbon monoxide. It is also available for
the formation and dissociation of ozone; we ignore this process because any
heat flux contributed in the lower atmosphere will not alter the number
density distribution greatly in the region of our interest.
The following reactions are considered here. The dissociation rate
coefficients a1 and a2 are derived by integrating the solar flux over all
the wavelengths applicable. The values of 7y and 7y were estimated from
the little infonrmation currently available on two-body reactions. Only
two-body association processes are considered, since the total atmospheric
number density in the dissociation region is too small for the corresponding
three-body processes to compete.
a=-12510 e- sec
CO + hv ->CO + 0 a = 1.25 x 10 7 e sec (1)
0+002 ~+CO2 ~ ~ iol8 3 -l()
0 + CO ->CO + 5.55 e.v. Y = 10l18 cm3 secl (2)
0+0 -02 +5.16 e.v. 2= 10 18 cm3 sec-1 (3)
02 + hv -O + 0 a2 = 2.5 x 106 e sec 1 (4)
where T is the total optical depth for the radiation considered. For this
purpose, the following mean cross section were used for each constituent:
(02) 1.62 x 10-18 cm2
a(C02) 2.16 x 1019 cm2
and T = a(02) j [ 2] dh + a(C02) [C] h
fh olh
The rate and continuity equations are then:
d [CO2 = (7 0 [00] [C] -a1 [C02]
d [(02 = 72g ( [02] (6)
dAC 2.9 2.4 2 [C(6)

5
[CO] = [0o + 2[02] (7)
CO2] = [CO2] + [co] (8)
where [CO ] is the initial number density of the primitive CO2 atmosphere.
Since we are now assuming CO2 is in a diffusive distribution, equation (8)
does not apply, i.e. [CO2] is known as a function of height.
For [CO2] = 0 = d [02] these equations can be combined to form
the equilibrium equations:
71 2 Y [2 = (C
- [0] + 2 —- [03 [CO] (9)
(21 = [l]2 (10o)
[Co] = [0] + 2 [02] (7)
For any given value of [CO002 one may merely solve the cubic equation
(9) for [0] and then calculate [02] and [CO].
The resulting distiribution is itself a first approximation. Before
a final distribution can be ascertained, we must obtain a temperature profile
for the upper atmosphere, since all the constituents will be in diffusive
equilibrium above a certain level.
IV. The Transition between Chemical and Diffusive Equilibrium
When the diffusive transport rate of an atmospheric constituent is
greater than its chemical reaction rate, the distribution of the constituent
is governed by diffusive, rather than chemical equilibrium. The transition
between the two processes takes place gradually near the lowest level at
which a molecule can diffuse upward for one scale height without suffering
a sufficient number of collisions to result in a reaction. This condition
occurs when the diffusion time constant td is comparable to the chemical
time constant t:
t D H2
td (11)
D

6
t c (< n)41 (12)
where H is the scale heights D the diffusion coefficient and (I n) the rate
coefficient for the reaction of the given species with another whose number
density is n. If we let:
D C T1/2
nT
where C is a constant and nT is the total number density, and use for H the
conventional expressi on:
kT
H =:mg
the criterion becomes:
n nT = C (f)2 T53/ (13)
The classical diffusion coefficient for a molecule of species 1 through
a gas of species 2 is
kT " m (.....)
- D _ kT | _ r 1 - (14)
12 mlV12 I' 8m1 Lm + J n"2 12
where m1 and m are the molecular masses, vl2 is the collision frequency of
species 1 with species 2 and c12 is the gas-kinetic cross section for
collision between the two species. The binary diffusion coefficient differs
from that derived by Chapman and Cowling (1960) in that the square bracket
in the radical is inverted and a slight difference in the constants. Values
in the International Critical tables give a T3/4 dependence rather than the
1/2
classical T/. Evaluation of the various expressions for diffusion coefficients for the temperatures of interest give results that are roughly equal,
For diffusion through a mixture of different gases, including self-diffusion,
we take the coefficient for species 1 as:
\T rm
Il1 -' S 1 m (15)
1 s
$

7
If j is the molecuiar weight of the diffusing species, and g is assumed
as 342 cm sec 2 for a height of abou-: 114.0 km in the Martian atmosphere,
10 T)S2 2
H- = 5.914. x 10 (T)2 cm2
The chemical time constants t for the pertinent reactions are as follows,
C
using the rate coefficients of reactions (1) to (4).
For CO the only loss mechanism is reaction (2). Thus
e (co) = To] sec
Reactions (2) and (3) provide loss mechanisms for 0. Then
018
t (o) = see
tc (() - [0] + [CO]. and 0.2 are 8 xr06 se
The photo-dissociation time constants for CO0 and 0 are 8 x 1 e sec
and 4 x 105 e" sec, respectively. With the assumed parameters, C02 will
nowhere be in chemical equilibrium. Above 110 km, all the constituents
will be in diffusive equilibrium.
One more condition must be imposed: the equation of continuity for
the atmosphere as a whole must be satisfied by the 0, 0 and CO distributions.
00 00 CO
O0 dh + 2 / 02 dh = CO dh
JJ 2- U
70 70 70
This neglects any escape of these constituents from the atmosphere.
At the bottom of the diffusion level then [CO] = ([0] + [O2])
V. Photo-ion Production
The rate of ion production at any level is found by multiplying the
number density of each constituent by its respective ionization rate q.
The rate of electron production is of course the sum of the ion production
rates. Estimation of the q's involves consideration of the solar spectrum
(Norton et al 1962) and the absorption cross sections (a). To facilitate
computation, certain wavelengths were grouped together, and the average
cross section and solar flux taken. Table I shows the values used. For

8
each 10 km level for the diffusive region above 110 km the optical depth
r(X) and e (%) were computed by summing a N H for each constituent, where
N is the number density and H the local scale height. Below 110 km the r's
were calculated from the density profile. Ionization cross sections were
taken from Nicolet et al (1960) and Norton et al (1962). Those for CO2
were assumed equal to those for CO for want of better information. Finally
the q's were obtained by summing J(X) cT(X) e'() over all the wavelengths
for each constituent, where J(x) is the photoionizing flux at the top of
the atmosphere in the given wavelength bands. Figure 3 shows the photoion
and electron production as a function of height. Two peaks occur in ion
production: a sharp one around 85 km due to the ionization of 02 and a
broad one around 140 km. The latter is due to the variation of cross
section and ionizing flux with wavelength and to the fact that the ionizing
flux at a given height depends on the total absorption above that height.
The sum of the peaks of each constituent produces a broad peak of electron
production. The prominant peak due to 0 ionization is largely attributable
2 o
to ionizing flux in the wavelength range from 1030 to 912 A which is not
absorbed by-the other constituents. A small 0 peak at 105 km is due to
0
radiation in the region 912 - 885 A which is not absorbed by either CO or
C02. It is interesting to note that the large 0 peak occurs despite the
relatively small amount of 02 in the atmosphere.
VI. Heat Balance
The heat produced by the ionization process is an important input to
the atmosphere and together with the heat produced by photodissociation
is a primary determinant of the temperature profile of the upper atmosphere.
Since we have already assumed a temperature profile to derive the various
distributions, we must now iterate the whole procedure until a selfconsistent temperature profile is achieved.
The equation of heat transport is:
00
AT1/2 2 = (E-R) dh (16)
h
1/2
where AT is the thermal conduction coefficient, E is the rate at which
heat is deposited by photochemistry at each level and R is the rate of

9
radiation at each level. The radiat:on teo m R includes radiation by each of
the important constituents as given in Bates 1951.
R(I) = 1,7 x 10~1 [ Oj exp (- 2) erg c3 sec (17)
R(CO2) =3.4 x 1028 [CO 2 n exp (- ) erg cm3 sec (8)
3.4 x 10- (co (18)
where n = total number density, and
R(CO) = 2.8 r 10o23 T2 [C] (19)
R(CO) represents rotational rather than vibrational radiation since the
latter is a much less efficient process.
-8 -1 -3 /2
The coefficient A is given as 180 erg cm sec deg 3 for O2 and
360 for 0 (Nicolet 1961); for CO it mlay be estimated as 180 erg cm't sec
eg3/2and for 02 as 130 erg cm& 1 -1-l3/2
deg s and for CO2 as 130 erg cm1 sec deg 3 As a mean value of A
at any level we adopted
130[C0ol + 180([O0] + [CO]) + 36010 (20)
A = -
i
Using the above equations (17) - (20) and the assumed value of T, we
may calculate dt which by integration gives a new T as a function of
height. The nature of the dependence of each R on T makes an analytic
solution impractical, thus necessitating an iterative procedure. An
exhaustive discussion of the radiation processes has been given by McElroy
et al 1965.
One of the uncertain quantities is the total heat input to the upper
atmosphere resulting from the photochemical processes. The value of Harris
and Priester 1962, was reduced for the distance of Mars, Our estimated
-2 -1
total of 1 erg cm sec was bivided between photoionization and photo-11
dissociation by attributing 1.9 x 10 ergs/photoion pair produced and
4.75 x 10']3 ergs/photodissociation.

10
VII. The Ionosphere
With this self-consistent temperature profile and the resulting consistent
constituents we are now ready to calculate the ionosphere by combining the
reactions of the neutral and ionized constituents. We begin by assuming
photochemical equilibrium,
Although only two chemical elements are present, the reactions are
about as complex as those in the lower terrestrial ionosphere. Table II
gives the reactions considered important and their estimated rate coefficients.
Recombination coefficients are called a; in each case except that for 0 + e
they represent dissociative recombination. Only two-body radiative recombination is important for 0, since the density of the atmosphere is too
low for a three-body process.
Some discussion of the assumed rate coefficients is in order. In most
cases they are not known to better than an order of magnitude, if that well.
The radictive recombination rate C1 is taken from Nawrocki and Papa (1961)
and the dissociative rate C2 from there and Norton et al (1962). a3 and a4
were estimated from Nawrocki and Papa (1961) and analogy to the NO + and
+ 2
NO rates. The atom-ion reaction is assumed to be mostly an ion-atom
exchange rather than a true charge exchange. The rate k4 is taken from
Norton et al (1962); the rate kg from Fjeldbo et al (1966). The other k's
are estimated largely by analogy, considering the energy balance and the
possibility of absorption of excess energy by excitation.
No three-body reactions are included because of the low total particle
density assumed; the probable rate of 10'27 cm6 sec- (Bortner 1965) makes
such reactions orders of magnitude less than two-body reactions in the
regions we are concerned with.
From the reactions in Table II, the following equilibrium equations
can be written using ne for the electron density:
(alne k[0] +) (1+ [CO + kkl[C2] [(] 2 + kO[g2][CO+) =21)
a2ne[021 = (q2 + k4O+] + k5[Co( + k(:CO]0 ) [O2] + k7[0][CO]
+ k9[CO2] [ 1 (22)
(c3n +k 1(0] + (k3 + k5) [02]) (CO] (q3 + ]) [CO] (23)

11
(4ne + (k2 + k7) [0] + k [02] + C O ) [O2 IC02] (24)
0+ {25)
n (0 [O] + [02 ] + [CO ] [c CO
Dissociative recombination turns out to be an unimportant mechanism for
removal o and C Ths makes solut of of equations (24) and (23)
2 +
very simple. Radiative recombination with 0 is also insignificant because
electrons and ions are in diffusion far below the level where the process
would manifest itself. The only direct recombination of importance is that of
02 +
Since Clne O 3rn and ajne can be neglected, CO02+ can be calculated
3 +
directly from equation (24) given the neutral abundances. Next [CO ] can
be calculated from equation (23), since we know [CO^ ]. Now we can calculate
[0 ] from equation (21).
We now can write [0, ]: n - [ CO - [CO2 ] and put this in
equation (22) which thereby becomes a quadratic equation in n. At this
point we can solve for [02] from equation (25).
The equilibrium equations were solved to obtain values of the number
densities of electrons and of the various ionic species as a function of
height, assuming chemical equilibrium. It turns out that CO and COm are
minor ions, the predominant ones being 0 and 0 2 is the predominate ion
below about 180 km, despite the fact that neutral 0 is more abundant than
neutral 0, This predominance is due to the multiplicity of charge exchange
reactions terminating with the ion 02 which has the lowest energy.
Thus far we have considered the ionosphere to be in chemical equilibrium.
A number of references (Bauer 1965; Hanson 1962, etc) have pointed out that
chemical equilibrium can obtain only if the diffusion rates are significant.y
smaller than the chemical reaction rates. There is a critical level above
which diffusive equilibrium predominates. Above 180 km the important
chemical reactions are those which remove the atomic oxygen ions. The
rate of this loss must be compared with the diffusion rate for 0 in an
atmosphere of predominantly 0 and CO.

12
The ambipolar diffusion coefficient c.'' in 0 or CO is given approximately
by
I i T 2 -1
D -= 7 x 10t cm sec
for the temperature we h1ivc assumed, where n is the total atmospheric number
density (Nawrocki and Papa 1961). The diffusion time constant
t = H2/D sec
d a
where H is now the Affective ion scale height. The chemical reaction rate is
given by the coefficient of [0 1 in equation (16). We find that t - td =
7 x 103 sec at 340 km. Using this information, tle electron-density profile
is faired in between the chemical lapse rate and diffusive rate. The
resulting ionosphere is shown in Figure 4.
Vt5I. iscussion
The ionosphere constructed by using the above procedures has three
peaks of comparable magnitude; in the lower two 0 is predominant while
the upper peak is a "Bradbury" peak of O. Study of the equilibrium ion
density equations reveals that with decreasing solar flux 0+ decreases much
more rapidly than 0 + This is largely due to the assumed high efficiency
+
of charge rearrangement between 0 and CO2. Although no study has yet been
made of a dynamic model, some broad conclusions can be drawn.
In times of low solar activity and for large solar zenith angles, the
top peak would disappear and the ionospheric temperature would decrease substantially} leavigi, a single thin but still rather dense 02 peak. The
density cf the single peak will depend on the latitude one is considering,
At low latitudes, the solar flux decreases steeply with time and a dynamic
analysis is necessary. For high latitudes however, the solar flux is never
very high, so that a much lower density would be expected. This is indeed
consistent with the Mariner IV data which measured the ionosphere -t; high
latitude as well as large solar zenith angle. If the Lemperature is as low
as 170~K, the 02 region would have a scale height of 5 km. It is obvious
that a temperature of 85 - 90~K which is hard to reconcil.e with heat balance
considerations is not necessary.

13
TA BLE 7I
0 0 Co
La J JJa C'* Jcr a Ja a C
A (x 10) (x 0o-9) (x o1017) (x 10o9) (x 10o17) (I 10-9) (x 10o17)
1030-912 115.8 59.56.515
912-885 54.0 39.3.728 14.5.269
885-770 38.4 39.8 1.017 11.7.305 49.5 1.290
705-200 127.9 217.0 1.700 112.2.880 157.8 1.234
200-130 23.0 15.2.661 7.6.330 13.13.571
130- 90 6.0 1.0.1667.52.0868.96.160
90- 44. 14 1.18.0838.588.04-17.923.0655
_o... 1.. o838.....

14
TABLE II
Reaction
0 + hv - 0 + e
CO + 0 - O + CO + 0.38 e.v.
CO++ 0-0+ + CO + 0.16 e.v.
CO + 02 0 + C02 + 0.73 e.v.
0 + e - 0 + hv
0 + + O2 + + 1.58 e.v.
02 + hv -> 02+ e
CO + 0 -0+ + C+ 1.93 e.v.
2 2
+ +
CO + 02 0 +2 0 + + 1.71 e.v.
CO + + + CO+ 1.36 e.v.
0 + e + 0 + 6.92 e.v.
CO + hv -CO0 + e
C02+ + CO - CO+ + C02 + 0.56 e.v.
+ +
CO + e — C + 0 + 2.85 e.v.
CO2 + hv - CO2 + e
O + C -0O + CO + 1.20 e.v.
CO2 + e - CO + 0 + 8.23 e.v.
Rate Coefficient
-1
l1
k1
k2
k3
1(X
k4
q2
as5
k6
k7
a2
q3
k8
"8
a3
q4
k9
a4
sec
= 2x
= x
= 2x
=2x
= 5 x
-1
sec
5 x
= 5 x
=5 x
= 2x
-1
sec
2 x
= 1x
-1
sec
= x
3 x
10-11.10
10
-1011
10
-12
10-11
o-11
10-7 <
-7
10
cm3
cm3
cm3
cm3
-1
sec
-1
sec
-1
sec
-1
sec
-1
sec
cm3 sec1
3 -1
cm3 sec
3 -1
cm3 sec
m3 s-1
cm sec
-1 3 -1
10" cm3 sec
-7 3 -1
10 cm sec
10~9 cm3 sec-1
10-7 cm3 sec'

T,.mp-erature (deg K)
20
Coo
4oo
300
0CO
100
0
0 300 4iC
~. ~ - - ~,.:.' -,D:,
I
I
I
I
I
I
J
i
I
I
i
I
I
I
I.og[ o 2].
T
/
l
" -, l — vib, re lax
I
<-. mescpause
0 2
0
12
LOG [ CO]
j-;iI I

200
h
(kin)
150
100
60
I L, i
i
i, —.,I:J
6_ 7
[02]
i~~
[o]
6 7
[0]
egos!
or.
I
m
9
~,.. w-w._.,^, I - a f
10 li
LOG'''i?'1;rE D.':i Rl (.3 )...J.2
12
13
14
15
FIGUlE 2

350
0;-, 0
300'a,5 o
h
I TJ m
150
100
-- -..."'
Aw I
\ e
CO
CO'
2
I
3
490
C0
/ p0
BliL-ll*~s5Llll;~I-~SOF=~CiC:rr Cllsryn-~~*~LUiOlJIL=;liYLEeTPl.rr.2x?
-,7 1:i=20 S0- --
5 1Q 20 50 lXll-/
II
I
I
I
_.,,-P _2=3.A >
i:~rC-::~Z~'T C~jJ~dll;~~?~~i~~ii ~ _r-~~Li;C;~~~lI~::_~jl ~ -?IWI~~jr-S L~-:~~L~Zr:L ~~:, C,- __U
0O2
0,5
1 2
200
500 1000 2'000
5000
10000
A I.ON 4 _' PR-DU' O -(c 3.e 1 3 )
2:.14 - - C.I-,'7 -'tUz! A ~~ cm'- D s c
T1 -r' 2..- -' i T 3
r? ^'i'j^-,

700
6oo
500
400oo
300
N
N<
N
\
N
N
N ah
\ll., -.
[0jj
[co+]
200
100
0
\'
4-. —l
l+]
Co J
N
CO'
2,~ _- - I
-- - _- i" __
104!O- i. ELECT Em AD'Xs, A I:. (m-3 )

References
1. Bates, D. R. (1951) "The Temperature of the Upper Atmosphere," Proc.
Phys. Soc. 64B, 805-820.
2. Bauer, S. J. (1965) "Hydrogen and Helium Ions," presented at IAGA
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