of 9
Chapter
I
Origin
and
early
history
Earth
is
the
namesake
of
the
terrestrial
planets,
also
known
as
the
inner
or
rocky
planets.
The
chemistry
of
meteorites
and
the
Sun
provide
constraints
on
the
composition
of
the
bulk
of
these
planets
and
they
provide
tests
of
theories
of
planetary
formation
and
evolution.
In
trying
to
understand
the
origin
and
structure
of
the
Earth,
one
can
take
the
geocentric
approach
or
the
ab
initio
approach.
In
the
former,
one
describes
the
Earth
and
attempts
to
work
backward
in
time.
For
the
latter,
one
attempts
to
track
the
evolu-
tion
of
the
solar
nebula
through
collapse,
cool-
ing,
condensation
and
accretion,
hoping
that
one
ends
up
with
something
resembling
the
Earth
and
other
planets.
Planets
started
hot
and
had
a
pre-history
that
cannot
be
ignored.
The
large-
scale
chemical
stratification
of
the
Earth
reflects
accretionary
processes.
Condensation
of
the
nebula
The
equilibrium
assemblage
of
solid
compounds
that
exists
in
a
system
of
solar
composition
depends
on
temperature
and
pressure
and,
there-
fore,
location
and
time.
The
condensation
behav-
ior
of
the
elements
is
given
in
Figures
1.1
and
1.2.
At
a
nominal
nebular
pressure
of
10
-
1
atm.
the
material
would
be
a
vapor
at
temperatures
greater
than
about
1900
K.
The
first
solids
to
condense
at
lower
temperature
or
higher
pres-
sure
are
the
refractory
metals
(such
as
W,
Re,
Ir
and
Os).
Below
about
1750
K
refractory
oxides
of
aluminum,
calcium,
magnesium
and
titanium
condense,
and
metallic
iron
condenses
near
1470
K
(Table
1.1
and
Figure
1.2).
Below
about
1000
K,
sodium
and
potassium
condense
as
feldspars,
and
a
portion
of
the
iron
is
stable
as
fayalite
and
ferrosilite
with
the
proportion
increasing
with
a
further
decrease
in
tempera-
ture.
FeS
condenses
below
about
750
K.
Hydrated
silicates
condense
below
about
300
K.
Differences
in
planetary
composition
may
depend
on
the
location
of
the
planet
,
the
location
and
width
of
its
feeding
zone
and
the
effects
of
other
planets
in
sweeping
up
material
or
perturb-
ing
the
orbits
of
planetesimals.
In
general,
one
would
expect
planets
closer
to
the
Sun
and
the
median
plane
of
the
nebula
to
be
more
refi·actory
rich
than
the
outer
planets.
On
the
other
hand
,
if
the
final
stages
of
accretion
involve
coalescence
of
large
objects
of
different
eccentricities,
then
there
n"lay
be
little
correspondence
between
bulk
chemistry
and
the
present
position
of
the
terres-
trial
planets
(Table
1.2).
There
is
evidence
that
the
most
refractory
elements
condensed
fi·om
the
solar
nebula
as
a
group,
unfractionated
from
one
another,
at
tem-
peratures
above
the
condensation
temperature
of
the
Mg-silicates.
Hence,
the
lithophile
refractory
elements
(Al,
Ca,
Ti,
Be,
Sc,
V,
Sr,
Y,
Zr,
Nb,
Ba,
rare-earth
elements,
Hf,
Ta,
Th
and
U
and,
to
some
extent,
Wand
Mo)
can
be
treated
together.
From
the
observed
abundance
in
samples
from
the
Moon,
Earth
and
achondrites,
there
is
strong
support
for
the
idea
that
these
e
lements
are
present
in
the
same
ratios
as
in
Cl
chondrites.
The
abundance
of
the
refi·actory
elements
in
a
given
planet
can
be
wealdy
constrained
from
4
ORIGIN
AND
EARLY
HISTORY
Be
Mg
Ca
Sc
Ti
v
Sr
y
Zr
Nb
Ba
La
Hf
Ta
W
Re
Os
lr
Ra
Ac
104
105
106
107
108
109
Uii
Ce
Pr
Nd
Pm
Sm
Eu
Gd
Tb
Dy
Ho
Er
Tm
Yb
Lu
Th
Pa
u
Np
Pu
Am
Cm
Bk
Cf
Es
Fm
Md
No
Lw
D
Early
Condensate
D
Silicate
D
Metal
I I
Volatiles
1300
- 600
K
lil:llll
.
"ll!
Volatiles
<600
K
Condensation
behavior
of
the
elements.
Short-lived
radioactive
elements
are
shown
in
italics
(after
Morgan
and
Anders,
1980)
.
g
Q)
5
~
Q)
(l_
E
~
0
Fraction
condensed
0 .2
0.4
Silicates
Fe
+
H
2
0
:;
FeO
+
H
2
Fe
+
H
2
S
:;
FeS
+
H2
Condensation
of
a
solar
gas
at
I
o-
4
atm
(after
d
Anders
,
I
980)
0.6
0 .8
(Na
,
K)AISi
3
0
8
MnS
Cu.
Ag,
Zn,
Ga
Ge
,
Sn
, Sb
F,
Cl
,
Br
,
I
1.0
(C)
Table
1. 1
I
Approximate
sequence
of
conden-
sation
of
phases
and
elements
from
a
gas
of
solar
composition
at
10
-
3
atm
total
pressure
Phase
Formula
Temperature
Hibonite
CaAI1
201
9
1770
K
Corundum
AI20
3
1758
K
Platinum
Pt,
W,
Mo,
Ta
metals
Zr,
REE,
U,
Th
Sc,
lr
Perovskite
CaTi0
3
1647
K
Melilite
Ca
2
AI
2
Si0
1-
Ca
2
Mg
2
Si
20 7
1625
K
Co
Spinel
MgAI
20 4
1513
K
AI
2
Si0
5
Metallic
iron
Fe,
Ni
1473
K
Diopside
CaMgSi
20
6
1450
K
Forsterite
Mg
2
Si0
4
1444
K
Anorthite
CaA
I
2
Si
2
0 a
1362
K
Ca
2
Si0
4
CaSi0
3
Enstatite
MgSi0
3
1349
K
Cr
20 3
P,
Au,
Li
MnSi0
3
MnS,
Ag
As,
Cu,
Ge
Feldspar
(Na,K)AISi
30 s
Ag,
Sb,
F.
Ge
Sn,
Zn,
Se,
Te,
Cd
Reaction
(Mg.Fe)2Si04
1000
K
products
(Mg,Fe)Si0
3
Troilite,
FeS,
(Fe,
Ni)S
700
K
pentlandite
Pb,
Bi,
In,
Tl
Magnetite
Fe
3
04
405
K
Hydrous
Mg
3
Si
2
0 7
2H
2
0.
etc.
minerals
Calcite
CaC0
3
<400
K
Ices
H2
0.
NH
3,
CH
4
<200
K
Anders
(1968),
Grossman
(1972),
Fuchs
and
others
(1973),
Gro
ss
man
a
nd
Larimer
(1974).
the
inferred
abundance
of
their
heat-producing
members.
uranium
and
thorium,
and
the
global
heat
flux
.
But
the
present
surface
heat
flow
does
not
accurately
represent
the
current
rate
of
heat
production.
A
large
fraction
of
the
present
heat
THEORIES
OF
PLANETARY
FORMATION
5
Table
1.2
I
Properties
of
the
terrestrial
planets
GM
R
p
D*
10
18
cm
3
/s
2
km
g/cm
3
I/ MR
2
km
Earth
398.60
6371
5.514
0.3308
14
Moon
4.903
1737
3.344
0.393
75
Mars
42.83
3390
3.934
0365
>28
Venus
324.86
6051
5.24
Mercury
22.0
2440
5.435
*Es
timated
crustal
thickness.
flow
is
due
to
cooling
of
the
Earth,
which
means
that
only
an
upper
bound
can
be
placed
on
the
uranium
and
thorium
content.
Nevertheless,
this
is
a
useful
constraint
particularly
when
com-
bined
with
the
lower
bound
on
potassium
pro-
vided
by
argon-40
and
estimates
of
K/U
and
Th
/U
provided
by
magmas
and
the
crust.
There
is
little
justification
for
assuming
that
the
volatile
ele-
ments
joined
the
planets
in
constant
propor-
tions.
In
this
context
the
volatiles
include
the
alkali
metals,
sulfur
and
so
forth
in
addition
to
the
gaseous
species.
Theories
of
planetary
formation
The
nature
and
evolution
of
the
solar
nebula
and
the
formation
of
the
planets
are
complex
subjects.
TI1e
fact
that
terrestrial
planets
did
in
fact
form
is
a
sufficient
motivation
to
keep
a
few
widely
dispersed
scientists
working
on
these
problems.
There
are
several
possible
mechanisms
of
planetary
growth.
Either
the
planets
were
assembled
from
smaller
bodies
(planetesimals),
a
piece
at
a
time,
or
diffuse
collections
of
these
bodies,
clouds
,
became
gravitationally
un
stab
le
and
collapsed
to
form
planetary-sized
objects
.
The
planets,
or
protoplanetary
nuclei.
could
have
formed
in
a
gas-free
environment
or
in
the
pres-
ence
of
a
lar
ge
amount
of
gas
that
was
subse-
quently
dissipated.
Some
hypotheses
speculate
that
l
arge
amounts
of
primordial
helium
dis-
solved
in
an
early
molten
Earth.
Others
assume
that
the
bulk
of
the
Earth
assembled
gas
-free
and
volatiles
were
brought
in
later.
TI1e
intermediate
6
ORIGIN
AND
EARLY
HISTORY
stages
of
planetary
assembly
involv
ed
impacts
of
lar
ge
objects.
The
final
stages
involved
sweeping
up
the
debris
and
collecting
an
outer
veneer
of
exotic
materials
from
the
Sun
and
the
outer
solar
system.
The
planets
originated
in
a
slowly
rotating
disk-shaped
'
solar
nebula'
of
gas
and
dust
with
solar
composition
.
The
temperature
and
pres-
sure
in
the
hydrogen-rich
disk
decreased
radially
from
its
center
and
outward
from
its
plane
.
The
disk
cooled
by
radiation,
mostly
in
the
direction
normal
to
the
plane,
and
part
of
the
incandes-
cent
gas
condensed
to
solid
'dust'
particles
.
As
the
particles
grew,
they
settled
to
the
median
pl
a
ne
by
collisions
with
particles
in
other
orbits,
by
viscous
gas
drag
and
gravitational
attraction
by
th
e disk.
The
total
gas
pressure
in
the
vicin-
ity
of
Earth's
orbit
may
have
been
of
the
order
of
10
-
1
to
10
-
4
of
the
present
atmospheric
pres-
sure.
The
particles
in
the
plane
formed
rings
and
gaps.
The
sedimentation
time
is
rapid,
but
the
processes
and
time
scales
involved
in
the
collec-
tion
of
small
objects
into
planetary-sized
objects
are
not
clear
.
Comets,
some
meteorites
and
some
small
satellites
may
be
left
over
from
the
early
stages
of
accretion
.
The
accretion-during-condensation
,
or
inho-
nl.ogeneous-accretion
,
hypothesis
leads
to
radi-
ally
zoned
planets
with
refractory
and
iron-rich
cores,
and
a
compositional
zoning
away
from
the
Sun
;
the
outer
planets
are
more
volatile-rich
because
they
form
in
a
colder
part
of
the
neb-
ula
.
Superimposed
on
this
effect
is
a
size
effect:
the
larger
planets,
having
a
larger
gravitational
cross
section,
collect
more
of
the
later
condens-
ing
(volatile)
material
but
they
also
involve
more
gravitational
heating.
In
the
widely
used
Safronov
cosmogo
ni-
cal
theo
r y
(1972)
it
is
assumed
that
the
Sun
ini-
tially
possessed
a
uniform
gas-dust
nebula.
The
nebula
evolves
into
a
torus
and
then
into
a disk.
Particles
with
different
eccentricities
and
incli-
nations
collide
and
settle
to
the
median
plane
within
a few
orbits.
As
the
disk
gets
denser,
it
breaks
up
into
many
dense
accumulations
where
th
e
self-gravitation
exceeds
the
disrupting
tidal
force
of
the
Sun.
As
dust
is
removed
from
the
bulk
of
the
nebula,
the
transparency
of
the
neb-
ula
increases,
and
a
large
temperature
gradient
is
established
.
If
the
relative
velocity
between
planetesimals
is
high
,
fragmentation
rather
than
accumula-
tion
will
dominate
and
planets
will
not
grow.
If
relative
velocities
are
low,
the
planetesimals
will
be
in
nearly
concentric
orbits
and
the
col-
lisions
required
for
growth
will
not
take
place
.
For
plausible
assumptions
regarding
dissipation
of
energy
in
collisions
and
size
distribution
of
the
bodies,
mutual
gravitation
causes
the
mean
relative
velocities
to
be
only
somewhat
less
than
the
escape
velocities
of
the
larger
bodies
.
Thus,
throughout
the
entire
course
of
planetary
growth
,
the
system
regenerates
itself
such
that
the
larger
bodies
would
always
grow
.
The
for-
mation
of
the
giant
planets,
however,
may
have
disrupted
planetary
accretion
in
the
inner
solar
system
and
the
asteroid
belt
.
TI1.e
initial
stage
in
the
formation
of
a
planet
is
the
condensation
in
the
cooling
nebula.
The
first
solids
appear
in
the
range
1750-1600
K
and
are
oxides,
silicates
and
titanates
of
calcium
and
aluminum
and
refractory
metals
such
as
the
plat-
inum
group.
These
minerals
(such
as
corundum
,
perovskite,
melilite)
and
elements
are
found
in
white
inclusions
(chondrules)
of
certain
mete-
orites,
most
notably
in
Type
III
carbonaceous
chondrites.
These
are
probably
the
oldest
surviv-
ing
objects
in
the
solar
system.
Metallic
iron
con-
denses
at
relatively
high
temperature
followed
shortly
by
the
bulk
of
the
silicate
material
as
forsterite
and
enstatite
.
FeS
and
hydrous
minerals
appear
at
very
low
temperature,
less
than
700
K.
Volatile-rich
carbonaceous
chondrites
have
for-
mation
temperatures
in
the
range
300-400
K,
and
at
least
part
of
the
Earth
must
have
accreted
from
material
that
condensed
at
these
low
tempera-
tures.
The
presence
of
He
,
C0
2
and
H
2
0
in
the
Earth
has
led
some
to
propose
that
the
Earth
is
made
up
almost
entirely
of
cold
carbonaceous
chondritic
material
-
the
cold-accr
etion
hypoth
e-
sis.
Even
in
some
current
geochemical
models,
the
lower
mantle
is
assumed
to
be
gas-rich,
and
is
speculated
to
contain
as
much
helium
as
the
ca
rbonaceous
chondrites
. This
is
unlikely
.
The
volatile-rich
material
may
have
come
in
as
a
lat
e ve
neer
-
the
inhomogenous
accretion
hypothesis.
Even
if
the
Earth
accreted
slowly,
compared
to
cooling
and
condensation
times,
the
later
stages
of
accretion
could
involve
material
that
condensed
further
out
in
the
nebula
and
was
later
perturbed
into
the
inner
solar
system.
A
drawn-out
accretion
time
does
not
imply
a
cold
initial
condition.
Large
impacts
reset
the
ther-
Inometer.
The
early
history
of
planets
was
a
very
vio-
lent
one;
collisions,
radioactive
heat
and
core
formation
provided
enough
energy
to
melt
the
planet.
Cooling
and
crystallization
of
the
planet
over
timescales
of
millions
of
years
resulted
in
its
chemical
differentiation
-
segregation
of
mate-
rial
according
to
density.
This
differentiation
left
most
of
the
Earth's
mantle
different
in
compo-
sition
from
that
part
of
the
mantle
from
which
volcanic
rocks
are
derived.
There
must
be
mate-
rial
that
is
complementary
in
composition
to
the
materials
sampled
by
volcanoes.
The
Earth
and
the
Moon
are
deficient
in
the
very
volatile
elements
that
make
up
the
bulk
of
the
Sun
and
the
outer
planets,
and
also
the
moderately
volatile
elements
such
as
sodium,
potassium,
rubidium
and
lead.
Mantle
rocks
con-
tain
some
primordial
noble
gas
isotopes.
(Reminder:
primordial
noble
gas
isotopes
is
a
Googlet.
If
it
is
typed
into
a search
engine
it
will
return
useful
information
on
the
topic,
including
defi-
nitions
and
references.
These
Googlets
will
be
sprinkled
throughout
the
text
to
provide
supplementary
infor-
mation.)
The
noble
gases
and
other
very
volatile
elements
were
most
likely
brought
in
after
the
bulk
of
the
Earth
accreted
and
cooled.
The
40
Ar
content
of
the
atmosphere
demonstrates
that
the
Earth
is
an
extensively
degassed
body;
the
atmo-
sphere
contains
about
70%
of
the
40
Ar
produced
by
the
decay
of
4
°K
over
the
whole
age
of
the
Earth.
This
may
imply
that
most
of
the
K
and
other
incompatible
elements
are
in
the
crust
and
shallow
mantle.
Magma
ocean
A
large
amount
of
gravitational
energy
is
released
as
particles
fall
onto
an
accreting
Earth,
enough
to
evaporate
the
Earth
back
into
space
as
fast
MAGMA
OCEAN
7
as
it
forms.
Even
small
objects
can
melt
if
they
collide
at
high
velocity.
The
mechanism
of
accre-
tion
and
its
time
scale
determine
the
fraction
of
the
heat
that
is
retained,
and
therefore
the
tem-
perature
and
heat
content
of
the
growing
Earth.
The
'initial'
temperature
of
the
Earth
was
likely
to
have
been
high
even
if
it
formed
from
cold
plan-
etesimals.
A
rapidly
growing
Earth
retains
more
of
the
gravitational
energy
of
accretion,
particu-
larly
if
there
are
large
impacts
that
can
bury
a
large
fraction
of
their
gravitational
energy.
Evi-
dence
for
early
and
widespread
melting
on
such
small
objects
as
the
Moon
and
various
meteorite
parent
bodies
attests
to
the
importance
of
high
initial
temperatures,
and
the
energy
of
accretion
of
the
Earth
is
more
than
15
times
greater
than
that
for
the
Moon.
The
intensely
cratered
surfaces
of
the
solid
planets
provide
abundant
testimony
of
the
importance
of
high-energy
impacts
in
the
later
stages
of
accretion.
During
accretion
there
is
a
balance
between
the
gravitational
energy
of
accretion,
the
energy
radiated
into
space
and
the
thermal
energy
pro-
duced
by
heating
of
the
body.
Latent
heats
asso-
ciated
with
melting
and
vaporization
are
also
involved
when
the
surface
temperature
gets
high
enough.
The
ability
of
the
growing
body
to
radi-
ate
away
part
of
the
heat
of
accretion
depends
on
how
much
of
the
incoming
material
remains
near
the
surface
and
how
rapidly
it
is
covered
or
buried
.
Devolatization
and
heating
associated
with
impact
generate
a
hot,
dense
atmosphere
that
serves
to
keep
the
surface
temperature
hot
and
to
trap
solar
radiation.
One
expects
the
early
stages
of
accretion
to
be
slow,
because
of
the
small
gravitational
cross
section
and
absence
of
atmosphere,
and
the
terminal
stages
to
be
slow,
because
the
particles
are
being
used
up.
The
tem-
perature
profile
resulting
from
this
growth
law
gives
a
planet
with
a
cold
interior,
a
tempera-
ture
peak
at
intermediate
depth,
and
a
cold
outer
layer
.
Superimposed
on
this
is
the
temperature
increase
with
depth
due
to
self-compression
and
possibly
higher
temperatures
of
the
early
accret-
ing
particles.
However,
large
late
impacts,
even
though
infrequent,
can
heat
and
melt
the
upper
mantle
.
Formation
of
99%
of
the
mass
of
Earth
probably
took
place
in
a few
tens
of
millions
of
8
ORIGIN
A
ND
EARLY
HISTORY
6
5
sz
"'
0
4
:::::.
Q)
':;
3
Q)
c.
E
2
~
Radius
(10
3
km)
Schematic
temperatures
as
a function
of
radius
at
three
stages
in
the
accretion
of
a
planet
(heavy
lines).
Temperatures
in
the
interior
are
initially
low
because
of
the
low
energy
of
accretion.
The
solidi
and
liquidi
and
the
melting
zone
in
the
upper
mantle
are
also
shown
.
Upper-mantle
melting
and
melt-solid
separation
is
likely
during
most
of
the
accretion
process
.
Silicate
melts,
enriched
in
incompatible
elements,
will
be
concentrated
toward
the
surface
throughout
accretion.
The
Earth,
and
perhaps
the
mantle,
will
be
stratified
by
intrinsic
density,
during
and
after
accretion
.
The
Melting
Zone
in
the
upper
mantle
or
a
near-surface
magma
ocean
processes
accreting
material.
Temperature
estimates
provided
by
D. Stevenson.
years
,
around
4.55
billion
years
ago
.
The
core
was
forming
during
accretion
and
was
already
in
place
by
its
end.
TI1ere
was
likely
not
a
co
r
e-
fo
rmin
g
eve
nt
.
Accretional
calculations,
taking
into
account
the
energy
partitioning
during
impact,
have
upper-mantle
temperatures
in
excess
of
the
melt-
ing
temperature
during
most
of
the
accretion
time
(Figure
1.3).
If
melting
gets
too
extensive
,
the
melt
moves
toward
the
surface
,
and
some
fraction
reaches
the
surface
and
radiates
away
its
heat.
A
hot
atmosphere.
a
thermal
bound-
ary
layer
and
the
presence
of
chemically
buoy-
ant
material
at
the
Earth's
surface
,
however
,
insu-
lates
most
of
the
interior
,
and
coolin
g
is
slow
.
Extensive
cooling
of
the
interior
can
only
occur
if
cold
surface
material
is
subducted
into
the
mantle.
TI1is
requires
a
very
cold
,
thick
thermal
boundary
layer
that
is
denser
than
the
under-
lying
mantle
.
This
plat
e tectonic
mod
e
of
mantle
convection
-
with
subduction
and
recycling
-
may
only
extend
back
into
Earth
history
about
1
Ga
(10
9
years
ago).
An
extensive
accumulation
of
basalt
or
olivine
near
the
Earth's
surface
durin
g
accretion
forms
a
buoyant
layer
that
resists
sub-
duction
.
An
extensively
molten
, slowly
cooling,
upper
mantle,
and
a
very
slowly
cooling
deeper
mantle
are
predicted
.
A
magma
ocean
freezes
from
the
bottom
but
a
thin
chill
layer
may
form
at
the
surface.
As
various
crystals
freeze
out
of
the
ocean
they
will
float
or
sink
,
depending
on
their
density
.
On
the
Moon,
plagioclase
crystals
float
when
they
freeze
and
this
is
one
explanation
of
the
anorthositic
highlands.
On
the
much
larger
Earth
,
the
aluminum
enters
dense
garnet
crys-
tals
and
a
deep
eclogite-rich
layer
is
the
result
.
Although
a
magma
ocean
may
be
convecting
vio-
lently
when
it
is
hot,
or
being
stirred
by
impacts,
at
some
point
it
cools
through
the
crystallization
temperatures
of
its
components
and
the
subse-
quent
gravitational
layering
depends
on
the
rela-
tive
cooling
rate
and
sinking
rates
of
the
crys-
tals.
Meanwhile,
new
material
is
being
added
from
space
and
is
processed
in
the
magma
ocean
.
A
chemically
stratified
Earth
is
the
end
result.
Accretional
and
convective
stirring
is
unlikely
to
dominate
over
gravitational
settling.
Magma
is
one
of
the
most
buoyant
products
of
mantle
differentiation
and
will
tend
to
stay
near
the
surface.
A
hot,
differentiated
planet
cools
by
the
heat-pipe
cool
in
g m
ec
h
anism
of
ma n
tle
convectio
n;
pipes
or
sheets
of
magma
remove
material
from
the
base
of
the
proto-crust
and
place
it
on
top
of
the
basaltic
pile,
which
gets
pushed
back
into
the
mantle,
cooling
the
interior
.
As
a
thick
basa
l t
crust
cools,
the
lower
portions
eventually
convert
to
dense
eclogite
-
instead
of
melting
-
and
delaminate.
TI1is
also
cools
off
the
interior
.
As
the
surface
layer
cools
further,
the
olivine-rich
and
eclogitic
parts
of
the
outer
layer
become
denser
than
th
e
interior
and
subduction
initiates.
At
this
point
portions
of
the
upper
mantle
are
rapidly
cooled
and
th
e
thermal
evolution
of
the
Earth
switches
ov
e r
to
the
plate-tectonic
era
.
Plate
tectonics
is
a
la te-sta
ge
method
for
coolin
g
off
the
interior
,
but
it
is
restricted
to
those
parts
of
the
interior
th
a t
are
less
dense
than
s
labs
. A
dense
primitiv
e
atmo-
sphere
and
buoyant
outer
layers
a
re
effectiv
e
insul
a
tors
and
serv
e
to
keep
the
crust
and
upper