Chapter
2
Comparative
planetology
The
sun,
with
all
those
planets
revolving
around
it
and
dependent
on
it,
can
still
ripen
a bunch
of
grapes
as
if
it
had
nothing
else
in
the
un1verse
to
do.
Go/ilea
Galilei
Before
the
advent
of
space
exploration
,
Earth
sci-
entists
had
a
handicap
almost
unique
in
science:
they
had
only
one
object
to
study.
Compare
thi
s
with
the
number
of
objects
available
to
ast
r
onomers,
particle
physicists,
biologists
and
sociologists.
Earth
theories
had
to
be
based
almost
entirely
on
evidence
from
Earth
itself.
Although
each
object
in
the
solar
system
is
unique,
we
have
learned
some
lessons
that
can
be
applied
to
Earth.
(1)
Study
of
the
Moon,
Mars
and
the
basaltic
achondrites
demonstrated
that
early
melting
is
ubiquitous.
(2)
Although
primitive
objects
,
such
as
the
car-
bonaceous
chondrites,
have
survived
for
the
age
of
the
solar
system,
there
is
no
evidence
for
the
survival
of
primitive
material
once
it
has
been
in
a
planet.
(3)
The
magma-ocean
concept
proved
useful
when
applied
to
the
Earth,
taking
into
account
the
differences
required
by
the
hi
g
her
pressures
on
the
Earth.
(4)
The
importance
of
great
impacts
in
the
early
history
of
the
planets
is
now
clear
.
(5)
Material
was
still
being
added
to
the
Earth
and
Moon
after
the
major
accretion
stage
and
the
giant
impacts,
and
is
still
being
added,
including
material
much
richer
in
the
noble
metals
and
noble
gases
than
occur
in
the
crust
or
mantle
.
(6)
TI1e
difference
in
composition
of
the
atmo-
spheres
of
the
terrestrial
planets
shows
that
the
original
volatile
compositions
,
the
extent
of
outgassing-
or
the
subsequent
processes
of
atmospheric
escape
-
have
been
quite
differ-
ent
.
(7)
We
now
know
that
plate
tectonics,
at
least
the
recycling
kind,
is
unique
to
Earth.
The
thick-
ness
and
average
temperature
of
the
litho-
sphere
and
the
role
of
phase
changes
in
basalt
seem
to
be
important.
Any
theory
of
plate
tec-
tonics
must
explain
why
the
other
terrestrial
planets
do
not
behave
like
Earth.
Although
the
Earth
is
a
unique
body,
and
is
the
largest
of
the
terrestrial
planets,
we
can
apply
lessons
learned
from
the
other
objects
in
the
solar
system
to
the
composition
and
evolution
of
the
Earth.
TI1e
Earth
is
also
an
average
terres-
trial
planet;
if
we
take
one
part
Mercury,
one
part
each
ofVenus
and
Mars,
and
throw
in
the
Moon
,
we
have
a
pretty
good
Earth,
right
size
and
den-
sity,
and
about
the
right
size
core
.
The
inner
solar
system
has
the
equivalent
of
two
Earths
.
Planetary
crusts
TI1e
total
crustal
volume
on
the
Earth
is
anoma-
lously
small,
compared
with
other
planets,
and
compared
with
its
crust-forming
potential,
but
it
nevertheless
contains
a
large
fraction
of
the
terrestrial
inventory
of
incompatible
elements.
The
thin
crust
on
Earth
can
be
explained
by
crustal
recycling
and
the
shallowness
of
the
basalt-eclogite
boundary
in
the
Earth.
Most
of
Earth's
'crust'
probably
resides
in
the
transition
region
of
the
mantle.
Estimates
of
bulk
Earth
chemistry
can
yield
a
basaltic
layer
of
about
10%
of
the
mass
of
the
mantle.
The
crust
of
the
Earth
is
enriched
in
Ca.
Al,
K
and
Na
in
comparison
to
the
mantle,
and
ionic-radii
considerations
and
experimental
petrology
suggest
that
the
crust
of
any
planet
will
be
enriched
in
these
constituents.
A
max-
imum
average
crustal
thickness
for
a
fully
dif-
ferentiated
chondritic
planet
can
be
obtained
by
removing
all
of
the
CaO,
with
the
available
Al
2
0
3
,
as
anorthite
to
the
surface.
This
operation
gives
a
crustal
thickness
of
about
100
kn1
for
Mars.
Incomplete
differentiation
and
retention
of
CaO
and
Al
2
0
3
in
the
mantle
will
reduce
this
value,
which
is
likely
to
be
the
absolute
upper
bound
(Earth's
crust
is
much
thinner
due
to
crustal
recycling,
delamination
and
the
basalt-eclogite
phase
change).
In
the
case
of
the
Earth.
up
to
60-70%
of
some
large-ion
elements
are
in
the
crust,
implying
that
about
30-40%
of
the
crustal
elements
are
in
the
mantle.
This
does
not
require
that
30-40%
of
the
mantle
is
still
in
a
primor-
dial
undegassed
state
as
some
geochemists
believe.
The
average
thickness
of
the
crust
of
the
Earth
is
only
15
km,
which
amounts
to
0.4%
of
the
mass
of
the
Earth.
The
crustal
thickness
is
5-10
km
under
oceans
and
30-50
km
under
older
continental
shields.
The
thickest
crust
on
Earth-
about
80
km
- is
under
young
actively
converg-
ing
mountain
belts.
The
parts
deeper
than
about
50
km
may
eventually
convert
to
eclogite,
and
fall
off.
The
situation
on
the
Earth
is
complicated,
since
new
crust
is
constantly
being
created
at
midoceanic
ridges
and
consumed
at
island
arcs.
The
continental
crust
loses
mass
by
erosion
and
by
delamination
of
the
lower
eclogitic
portions.
Continental
crust
is
recycled
but
its
total
volume
is
roughly
constant
with
time
.
Both
the
Moon
and
Mars
have
crustal
thicknesses
greater
than
that
of
the
Earth
in
spite
of
their
much
smaller
sizes,
and
probable
less
efficient
differentiation.
MERCURY-
FIRST
ROCK
FROM
THE
SUN
13
Mercury-
first
r
ock
from
t he
Sun
Mercury
is
5.5%
of
the
mass
of
the
Earth,
but
it
has
a
very
similar
density,
5.43
gfcm
3
.
Its
radius
is
2444
km.
Any
plausible
bulk
composi-
tion
is
about
60%
iron
and
this
iron
must
be
largely
differentiated
into
a
core.
Mercury
has
a
perceptible
magnetic
field,
appreciably
more
than
either
Venus
or
Mars,
probably
implying
that
the
core
is
molten.
Mercury's
surface
is
pre-
dominantly
silicate,
but
apparently
not
basaltic.
A
further
inference
is
that
the
iron
core
existed
early
in
its
history;
a
late
core-formation
event
would
have
resulted
in
a
significant
expansion
of
Mercury.
Mercury's
shape
may
have
significantly
changed
over
the
history
of
the
planet.
Tidal
de-
spinning
results
in
a less
oblate
planet
and
com-
pressional
tectonics
in
the
equatorial
regions.
Cooling
and
formation
of
a
core
cause
a
change
in
the
mean
density
and
radius.
A
widespread
system
of
arcuate
scarps
on
Mercury,
which
appear
to
be
thrust
faults,
provides
evidence
for
compressional
stresses
in
the
crust.
The
absence
of
normal
faults
suggests
that
Mercury
has
contracted.
This
is
evidence
for
cooling
of
the
interior.
One
factor
affecting
the
bulk
composition
of
Mercury
is
the
probable
high
temperature
in
its
zone
of
the
solar
nebula;
it
may
have
formed
from
predominantly
high-temperature
condensates.
If
the
temperature
was
held
around
1300
K
until
most
of
the
uncondensed
mate-
rial
was
blown
away,
then
a
composition
satis-
fYing
Mercury's
mean
density
can
be
obtained,
since
most
of
the
iron
will
be
condensed,
but
only
a
minor
part
of
the
magnesian
silicates.
Since
the
band
of
temperatures
at
which
this
condition
prevails
is
quite
narrow,
other
fac-
tors
must
be
considered.
Two
of
these
are
(1)
dynamical
interaction
among
the
material
in
the
terrestrial
planet
zones,
leading
to
compo-
sitional
mixing,
and
(2)
collisional
differentia-
tion.
A
large
impact
after
core
formation
may
have
blasted
away
much
of
the
silicate
crust
and
mantle.
Our
Moon
may
have
been
the
result
of
such
an
impact
on
proto-Earth.
On
Mars,
the
crust
is
locally
thinner
under
the
large
impact
basins.
14
COMPARATIVE
PLANETOLOGY
Terrestrial
bodies
were
subjected
to
a
high
flux
of
impacting
objects
in
early
planetary
his-
tory.
The
high-flux
period
can
be
dated
from
lunar
studies
at
about
3.8
billion
years
ago.
The
large
basins
on
the
surface
of
Mercury
formed
during
this
period
of
high
bombardment
.
Later
cooling
and
contraction
apparently
were
respon-
sible
for
global
compression
of
the
outer
surface
and
may
have
shut
off
volcanism.
On
the
Earth,
volcanism
is
apparently
restricted
to
the
extend-
ing
regions.
Venus
Venus
is
320
km
smaller
in
radius
than
the
Earth
and
is
about
4 .
9%
less
dense.
Most
of
the
dif-
ference
in
density
is
due
to
the
lower
pressure,
giving
a
smaller
amount
of
self-compression
and
deeper
phase
changes.
Venus
is a
smoother
planet
than
the
Earth
but
has
a
measurable
triaxiality
of
figure
and
a 0.34
km
offset
of
the
center
of
the
figure
from
the
center
of
mass.
This
offset
is
much
smaller
than
those
of
the
Moon
(2
Ian),
Mars
(2.5
km)
and
Earth
(2
.1
km).
In
contrast
to
the
bimodal
distribution
of
Earth's
topography,
representing
continent-ocean
differences,
Venus
has
a
narrow
unimodal
height
distribution
with
60%
of
the
surface
lying
within
500
m
of
the
mean
elevation.
This
difference
is
probably
related
to
erosion
and
isostatic
differ-
ences
caused
by
the
presence
of
an
ocean
on
Earth.
For
both
Earth
and
Venus
the
topogra-
phy
is
dominated
by
long-wavelength
features.
Most
of
the
surface
of
Venus
is
gently
rolling
ter-
rain.
The
gravity
and
topography
are
positively
correlated
at
all
wavelengths.
On
Earth
most
of
the
long-wavelength
geoid
is
uncorrelated
with
surface
topography
and
is
due
to
deep-mantle
dynamics
or
density
variations.
The
other
respects
in
which
Venus
differs
markedly
from
the
Earth
are
its
slow
rotation
rate,
the
absence
of
a
satellite,
the
virtual
absence
of
a
magnetic
field,
the
low
abundance
of
water,
the
abundance
of
primordial
argon,
the
high
sur-
face
temperature
and
the
lack
of
obvious
signs
of
subduction.
From
crater
counts
it
appears
that
the
age
of
the
surface
of
Venus
is
300-500
million
years
old,
much
less
than
parts
of
the
Earth's
surface
.
The
oceanic
crust
on
the
Earth
is
renewed
every
200
million
years
but
the
conti-
nents
survive
much
longer.
If
Venus
had
an
identical
bulk
composition
and
structure
to
the
Earth,
then
its
mean
den-
sity
would
be
about
5.34
gfcm
3
.
By
'identical
structure'
I
mean
that
(1)
most
of
the
iron
is
in
the
core,
(2)
the
crust
is
about
0.4%
of
the
total
mass
and
(3)
the
deep
temperature
gradi-
ent
is
adiabatic
(an
assumption).
The
high
sur-
face
temperature
of
Venus,
about
740
K,
would
have
several
effects;
it
would
reduce
the
depth
at
which
the
convectively
controlled
gradient
is
attained,
it
would
deepen
temperature-sensitive
phase
changes
and
it
may
prevent
mantle
cool-
ing
by
subduction
.
The
density
of
Venus
is
1.2-1.9%
less
than
that
of
the
Earth
after
correcting
for
the
dif-
ference
in
pressure.
This
may
be
due
to
differ-
ences
in
iron
content,
sulfur
content,
oxidization
state
and
deepening
of
the
basalt-eclogite
phase
change
. Most
of
the
original
basaltic
crust
of
the
Earth
subducted
or
delaminated
when
the
upper-
mantle
temperatures
cooled
into
the
eclogite
sta-
bility
field
.
The
density
difference
between
basalt
and
eclogite
is
about
15%.
Because
of
the
high
surface
temperature
on
Venus,
the
upper-mantle
temperatures
are
likely
to
be
200-400
K
hotter
in
the
outer
300
km
or
so
than
at
equivalent
depths
on
Earth,
or
melting
is
more
extensive
.
This
has
interesting
implications
for
the
phase
relations
in
the
upper
mantle
and
the
evolution
of
the
planet.
In
particular,
partial
melting
in
the
upper
mantle
would
be
much
more
extensive
than
is
the
case
for
the
Earth
except
for
the
fact
that
Venus
is
probably
deficient
in
the
volatile
and
low-molecular-weight
elements
that
also
serve
to
decrease
the
melting
point
and
viscosity.
Crust
can
be
much
thicker
because
of
the
deepening
of
the
basalt-eclogite
phase
boundary
.
Schematic
geotherms
are
shown
in
Figure
2.1
for
surface
temperatures
appropriate
for
Earth
and
Venus.
With
the
phase
diagram
shown,
the
high-temperature
geotherm
crosses
the
solidus
at
about
85
lan.
With
other
plausible
phase
rela-
tions
the
eclogite
field
is
entered
at
a
depth
of
about
138
km.
For
Venus,
the
lower
gravity
and
outer-layer
densities
increase
these
depths
by
about
20%;
thus,
we
expect
a
surface
layer
1500
~
Q.l
':;
~
1000
Q.l
0.
E
~
Q.l
u
"'
'1::
::>
(f)
500
300
Depth
(km)
Schematic
geotherms
for
the
Earth
with
different
surface
temperatures
.
Note
that
the
eclogite
stability
field
is
deeper
for
the
higher
geotherms
and
that
a
partial
melt
field
intervenes
between
the
basaltic
crust
and
the
rest
of
the
upper
mantle.
Basaltic
material
in
the
eclogite
field
will
sink
through
much
of
the
upper
mantle
and
will
be
replaced
by
peridotite.
Shallow
subduction
of
basaltic
crust
leads
to
remelting
in
the
case
of
Venus
and
the
early
Earth.
Conversion
to
eclogite
leads
to
lower
crustal
delamination
or
deep
subduction
for
the
present
Earth.
This
figure
explains
why
the
low-density
crust
can
be
no
thicker
than
about
SO
km
on
Earth.
The
depth
scale
is
for
an
Earth-size
planet
with
the
colder
geotherm
and
present
crust
and
upper-mantle
densities.
For
Venus,
with
smaller
gravity,
higher
temperatures
and
low-density
crust
replacing
part
of
the
upper
mantle,
the
depths
are
increased
by
about
20%
.
of
100-170
km
thickness
on
Venus
composed
of
basalt
and
partial
melt.
On
the
present
Earth,
the
eclogite
stability
field
is
entered
at
a
depth
of
40-60
km.
If
the
interior
of
Venus
is
dry
it
will
be
stronger
at
a
given
temperature
and
will
have
a
higher
solidus
temperature.
A
large
amount
of
basalt
has
been
produced
by
the
Earth's
mantle,
but
only
a
thin
veneer
is
at
the
surface
at
any
given
time.
There
must
there-
fore
be
a
substantial
amount
of
eclogite
in
the
mantle,
the
equivalent
of
about
200
krn
in
thick-
ness.
If
this
were
still
at
the
surface
as
basalt,
the
Earth
would
be
several
percent
less
dense.
Correcting
for
the
difference
in
temperature,
MARS
IS
surface
gravity
and
mass
and
assuming
that
Venus
is
as
well
differentiated
as
Earth,
only
a
fraction
of
the
basalt
in
Venus
would
have
con-
verted
to
eclogite.
This
would
make
the
uncom-
pressed
density
of
Venus
about
1.5%
less
than
Earth's
without
invoking
any
differences
in
com-
position
or
oxidation
state.
Thus,
Venus
may
be
close
to
Earth
in
composition.
It
is
possible
that
the
present
tectonic
style
on
Venus
is
similar
to
that
of
Earth
in
the
Archean,
when
temperatures
and
temperature
gradients
were
higher
.
If
the
Moon,
Mercury
and
Mars
have
molten
iron
core,
it
is
probable
that
Venus
does
as
well.
The
youthful
age
of
the
surface
of
Venus
has
been
attributed
to
a
gl
o
bal
res
urfa
c i
ng
ev
en
t .
The
cratering
record
indicates
that
the
global
resurfacing
event
,
about
300
my
ago,
was
followed
by
a
reduction
of
volcanism
and
tec-
tonism.
Delamination
of
thick
basaltic
crust,
foundering
of
a
cold
thermal
boundary
layer
and
a
massive
reorganization
of
mantle
convection
are
candidates
for
the
resurfacing
event.
Mantle
convection
itself
is
strongly
controlled
by
surface
processes
and
changes
in
these
processes.
Mars
Mars
is
about
one-tenth
of
the
mass
of
Earth
.
The
uncompressed
density
is
substantially
lower
than
that
of
Earth
or
Venus
and
is
very
similar
to
the
inferred
density
of
a
fully
oxidized
(minus
C
and
H
2
0)
chondritic
meteorite
.
The
moment
of
inertia,
however
,
requires
an
increase
in
den-
sity
with
depth
over
and
above
that
due
to
self-
compression
and
phase
changes
,
indicating
the
presence
of
a
dense
core.
This
in
turn
indicates
that
Mars
is
a
differentiated
planet.
The
tenuous
atmosphere
of
Mars
suggests
that
it
either
is
more
depleted
in
volatiles
or
has
experienced
less
outgassing
than
Earth
or
Venus.
It
could
also
have
lost
much
of
its
early
atmo-
sphere
by
large
impacts
.
Geological
evidence
for
running
water
on
the
surface
of
Mars
suggests
that
a
large
amount
of
water
is
tied
up
in
per-
mafrost
and
ground
water
and
in
the
polar
caps.
Whether
Mars
had
standing
water-
oceans
and
lakes
-
for
long
periods
of
time
is
currently
being
debated.
The
high
4 0
Ar
f
36
Ar
ratio
on
Mars
,
16
COMPARATIVE
PLANETOLOGY
ten
times
the
terrestrial
value,
suggests
either
a
high
potassium-40
content
plus
efficient
outgas-
sing,
or
a
net
depletion
of
argon-36
and,
possibly,
other
volatiles.
If
Mars
is
volatile-rich,
compared
to
Earth,
it
should
have
more
K
and
hence
more
argon-40.
Early
outgassed
argon-36
could
also
have
been
removed
from
the
planet.
SNC
meteorites
have
trapped
rare-gas
and
nitro
ge
n
contents
that
differ
from
other
mete-
orites
but
closely
match
those
in
the
martian
atmosphere.
If
SNC
meteorites
come
from
Mars,
then
a
relatively
volatile-rich
planet
is
implied,
and
the
atmospheric
evidence
for
a
low
volatile
content
for
Mars
would
have
to
be
rationalized
by
the
loss
of
the
early
accretional
atmosphere.
Mars
is
more
susceptible
to
atmospheric
escape
than
Venus
or
Earth
owing
to
its
low
g
ravity
.
The
surface
of
Mars
appears
to
be
weathered
basalt.
The
dark
materials
at
the
s
urface
contain
basaltic
minerals
and
hematite
and
sulfur-rich
material
,
and
there
is
evidence
for
the
past
action
of
liquid
water.
The
large
volcanoes
on
Mars
are
similar
in
form
to
shield
volcanoes
on
Earth.
Andesite
- a
possible
indicator
of
pl
ate
tectonics
-
has
been
proposed
as
a
component
of
martian
soil
but
this
is
controversial;
weathered
basalt
can
explain
the
ava
ilable
data.
The
topography
and
gravity
field
of
Mars
indi
-
cate
that
parts
of
Mars
are
grossly
out
of
hydro-
static
equilibrium
and
that
the
crust
is
highly
variable
in
thickness.
If
variations
in
the
g rav-
ity
field
are
attributed
to
variations
in
crustal
thickness,
reasonable
values
of
the
density
con-
trast
imply
that
the
average
crustal
thickness
is
at
least
45
km,
and
the
maximum
crustal
thick-
ness
may
reach
100
km.
Giant
impacts
may
have
removed
most
of
the
crust
beneath
the
basins,
replacing
crustal
material
by
uplifted
mantle.
If
so,
the
crust
was
in
place
in
early
martian
his-
tory,
consistent
with
other
evidence
throughout
the
solar
system
for
rapid
early
planetary
differ-
entiation.
On
Earth,
delamination
of
lower
crust
produces
a
thinning
but
the
whole
crust
is
not
involved.
The
only
direct
evidence
concerning
the
inter-
nal
structure
of
Mars
is
the
mean
density,
moment
of
inertia,
topography
and
gravity
field
.
The
mean
density
of
Mars,
corrected
for
pressure
,
is
less
than
that
of
Earth,
Venus
and
Mercury
MARS
Pm
-
3 .
54
g/
cm
3
Core
radius
(R
c
I
R )
Radius
of
the
core
versus
density
of
core
for
Mars
models
.
The
points
are
for
meteorites
with
all
of
the
FeS
and
free
iron
and
nickel
differentiated
into
the
core.
The
dashed
line
shows
how
core
density
is
related
to
core
s ize
in
the
Fe-FeS
system.
but
greater
than
that
of
the
Moon.
This
impli
es
either
that
Mars
has
a
small
total
Fe-Ni
content
or
that
the
FeO/Fe
ratio
varies
among
the
plan-
ets.
Plausible
models
for
Mars
can
be
constructed
that
have
solar
or
chondritic
values
for
iron,
if
most
or
all
of
it
is
taken
to
be
oxidized.
With
such
broad
chemical
constraints,
mean
density
and
moment
of
inertia
and
under
the
ass
ump-
tion
of
a
differentiated
planet,
it
is
possible
to
trade
off
the
size
and
density
of
the
core
and
density
of
the
mantle.
The
mantle
of
Mars
is
presumably
composed
mainly
of
silicates,
which
can
be
expected
to
undergo
one
or
two
major
phase
changes,
each
involving
a
10
%
increase
in
density.
To
a
good
approximation,
these
phase
changes
will
occur
at
one-third
and
two-thirds
of
the
radius
of
Mars.
The
deeper
phase
change
will
not
occur
if
the
radius
of
the
core
exceeds
one-third
of
the
radius
of
the
planet.
The
curve
in
Figure
2.2
is
the
locus
of
possible
Mars
models
. Clearly,
the
data
can
accommodate
a
small
dense
core
or
a
large
light
core.
The
upper
limit
to
the
density
of
the
core
is
probably
close
to
the
density
of
iron
,
in
which
case
the
core