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# Earth’s Core

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E
Earths Core
William F. McDonough
Department of Geology, University of Maryland, College
Park, MD, USA
Definition
The composition of the Earths core is essentially 85% Fe, 5%
Ni, and ~10% other, by weight. The details are more compli-
cated, with the otherbeing a light element component,
which results in the core having on average a lower mean
atomic number than that of iron. Birch (1952,1964) recog-
nized a 10% difference in the density of the outer core relative
to an iron-nickel alloy at core pressure and temperature con-
ditions. Anderson and Isaak (2002) recommend that the core
density decit is between 3% and 7%. Both the liquid outer
core (~95% by mass) and solid inner core (~5%) contain a
fraction of a light element (nominally in a ~2/1 proportion,
respectively).
Properties
For more than a century the Earth has been described as
having a metallic core surrounded by a silicate shell. A brief
history of this topic is given in McDonough (2014). Emil
Wiechert in 1897 rst subdivided the Earths interior into
these two main components and later, his graduate student
Beno Gutenberg, using seismological data, established the
core-mantle boundary (CMB) at 2900 km depth, which is
within uncertainty the accepted depth for the boundary.
First-order physical aspects of the core are listed in Table 1
and are based on fundamental constraints derived from seis-
mology, geodetics of the planets, mineral physics, and exper-
imental studies, with limited additional constraints from
geochemistry and cosmochemistry. Seismology reveals that
the CMB is more than half an Earths radius. The outer core
(OC) is liquid (absence of a shear-wave) and the inner core
(IC) is solid (given observation of PKJKP waves (i.e.,
P-waves in mantle and OC that converts to an S-wave in IC
then back to P-wave as it leaves the IC). Geodesy documents
the coefcient of the planets moment of inertia (i.e., a dense
core surrounds a less dense silicate shell) and libration con-
rms the liquid state of the outer core. Thermodynamic anal-
ysis and experimental studies document the conditions of
melting and crystallization of Fe-Ni alloys at core conditions
and set the upper temperature limits at the inner-outer core
boundary. Meteorites conrm that Fe and Ni are chemical
partners and their ratios in primitive and iron meteorites
inform us of their initial proportions in the Earth.
The maximum in the Earths gravitational acceleration
occurs at just above the CMB, where being close to the high
density core has its greatest effect on the gravity eld. Con-
sequently, g(i.e., acceleration of gravity) increases in
roughly linear fashion from center to CMB (i.e., approxi-
mately a uniform density sphere) and then remains between
9.8 and 10.5 m/s
2
from the surface to the CMB. The Earths
dynamic gure for the gravitational eld is known to six
signicant gures and in combination with seismological
constraints indicates that the excess ellipticity of the CMB is
slight (~0.2 km) and that the CMB surface topography is
smooth at long-wavelength being on the order of 13 km of
peak-to-peak topographic amplitude (Table 1).
Formation
The cores origin dates back to the formation of the planet,
with planet accretion occurring via accumulation and agglom-
eration of increasing size of particles and later planetesimals.
Our solar system owes it origin to a supernova (or multiple
supernovae) where the mass and momenta of the supernova
combined with a fragment of an interstellar gas-dust cloud to
rotate, collapse, condense, and centrally concentrate mass,
#Springer International Publishing AG 2017
W.M. White (ed.), Encyclopedia of Geochemistry,
DOI 10.1007/978-3-319-39193-9_258-1
internally controlled by a balance of gravitational and cen-
tripetal forces. Because of its mass the Sun, and to a lesser
extent Jupiter, dominated accretion and condensation of the
inner rocky planets, which formed relatively early (a few
10
6
years) and grew by the accretion of material owing
into and less so out of their respective feeding zone (rst
few AU surrounding the Sun).
Growth by accretion of a planet does not require differen-
tiation of the body into a core and mantle, although differen-
tiation likely occurred in stages during growth and included
accretion of predifferentiated planetismals. For the Earth, it is
likely that as accretion progressed the proto-Earth at some
point reach melting temperatures for metallic liquids (lower
than that for silicate liquids) and because of the large density
contrasts (i.e., silicates being ~3000 and iron metal being
~7900 kg/m
3
at surface conditions) gravitational separation
of a metallic core quickly ensued. The mean time of formation
of the bulk of the Earths core is, however, an open question
with best estimates being between 30 and 60 Ma (million
years) after the formation of the solar system (an age
established by the oldest dated materials in the solar system,
Hf-W isotope systematics of iron meteorites, and the age of
core formation can be considered as the mean age (or mean
lifetime), which is a scaling time reecting when much of the
mass of metallic atoms were separated. The initial mass
associated with core formation was likely signicant and
rapidly separated. Subsequent additions to the core might
have followed some exponentially decreasing mass versus
time function.
Earths Core, Table 1 Physical properties of the Earths core
Mass Value (!) Units Ref (% Planet)
Earth 5.97218(60) "10
24
kg 1 100%
Inner core 9.68 "10
22
kg 2 1.6%
Outer core 1.84 "10
24
kg 2 30.7%
Core 1.93 "10
24
kg 2 32.3%
Mantle 4.015 "10
24
kg 1,2,3 67.7%
Crust (ocean + continental) 2.73(48) "10
22
kg 3 0.46%
Inner core to core (%) 5.0%
Inner-outer core boundary 1220 !10 km 4 19%
Core-mantle boundary 3483 !5 km 4 55%
Mean radius of the Earth 6371.23 (0.01) km 1 100%
Volume
Inner core 7.61 "10
9
km
3
0.7%
Outer core 1.69 "10
11
km
3
16%
Bulk core 1.77 "10
11
km
3
16%
Silicate Earth 9.14 "10
11
km
3
84%
Earth 1.083 "10
12
km
3
100%
Density
Outermost inner core (solid) 12,830 kg/m
3
5
Innermost outer core (liquid) 12,010 kg/m
3
5
Inner-outer core Ddensity 820 (180) kg/m
3
5
Average outer core density 11,160 (60) kg/m
3
5
Average inner core density 13,070 (260) kg/m
3
5
CMB (Core-mantle boundary)
Peak-to-peak topography <3 km 6
CMB ellipticity ~0.2 km 6
Thermal and pressure data
Top of outer core 4000 !500 K 7,8 137 GPa
Top of inner core 5700 !550 K 7,8 330 GPa
Heat ow across CMB 515 TW 7
Moment of inertia constants
Equatorial moment of inertia (I) 0.3299765 Ma
2
2
Mean moment of inertia (I) 0.330690 (9) MR
2
0
1
1
Chambat et al. (2010);
2
Yoder (1995);
3
Huang et al. (2013);
4
Masters and Shearer (1995);
5
Masters and Gubbins (2003);
6
Sze and van der
Hilst (2003);
7
Tsuchiya et al. (2016);
8
Fischer (2016)
2 Earths Core
A second constraint on the time of core formation comes
from understanding the origin of the Moon and the dating of
its oldest rocks. This constraint is based on the fact that the
Moon possesses a bulk composition that approximates that of
the Earths mantle (i.e., they share the same oxygen isotopic
composition to six signicant gures), with a bit more iron in
the Moon (Young et al. 2016). The Moons origin is often
ascribed to being due to a giant impact (collision) between a
nearly completely formed Earth and a Mars sized impactor
(Mars is ~10% by mass the size of the Earth, whereas the
Moons mass is 7.346 "10
22
kg, which is 1.2% of the mass of
the Earth) (Canup 2012; Cuk and Stewart 2012). Alternative
scenarios for the moon-forming event(s) allow for a series of
lesser, but still substantial, impacts that also satisfy the obser-
vational constraints (Cuk et al. 2016). The material ejected
during this event coalesced into a corotating accretion disk
around the Earth, possessed only a small fraction of metal
from the Earths core or the impactor, and subsequently
accreted to form the Moon. The general features of this
model (Canup 2012; Cuk and Stewart 2012; Cuk et al.
2016) are consistent with hydrodynamic models of accretion
that seek to reconcile the angular momentum of the Earth-
Moon system and isotopic constraints.
Major questions about the composition of the Earths core,
its age, the age of the inner core, present conditions, and
dynamics of its formation remain unresolved. Debate focuses
on the following questions for which we have limited obser-
vational constraints:
Formation Conditions and Composition
What was the integrated composition of material feeding
the Earth (i.e., chondrites)?
What is the temperature and pressure evolution of planet
growth and core formations?
What was the integrated fO
2
, fH
2
, and fS
2
(gas fugacity) of
core formation?
How did the redox state of core formation evolve (early to
late evolution)?
Was segregation and sinking of metal from silicate an
equilibrium process throughout the mantle or only for a
portion thereof?
Was core separation accomplished in the presence a
magma ocean or not?
Did the process of giant impact collision coalesce and
emulsify the cores of the Earth and its impactor and
re-establish metal-silicate equilibrium?
What is the duration and timing of core formation (e.g.,
182
Hf-
182
W. isotope system)?
How much S, C, N, and H (i.e., volatile elements) are in
core?
How much O, Si, and Mg (i.e., major elements) are in
core?
Are there radioactive elements in the core and if so, how
much?
Inner-Outer Core
When did the inner core start to form?
What is the crystal structure of the inner core?
How much and what is (are) the light element(s) in the
inner core?
How much and what is (are) the light element(s) in the
outer core?
Does enrichment of buoyant, light-element-rich liquids,
produced at the inner-outer core boundary, drive compo-
sitional convection and the geodynamo?
To what degree does crystallization of the inner core drive
compositional convection?
Many unknowns remain and detailed constraints on the cores
composition come almost exclusively from the chemical and
isotopic composition of the silicate Earth and meteorites.
Building Compositional Models of the Planet
and Core
Our view of planetary compositional modeling is explicitly
informed by the compositions of chondrites, the most primi-
tive rocks of the solar system. Chondritic meteorites, which
make up the bulk of the meteoritic material (~95% of the
falls), fell to the Earth during historical times and thus shape
our view of the building blocks of the planets (Engel and
McDonough 2016). CI1 carbonaceous chondrites have the
most primitive composition of the various chondrites and
closely match the spectroscopically derived composition of
the solar photosphere for all but H, C, N, O, and the noble
gases (McDonough 2016). [Note, that the chondrite type is CI
and the metamorphic grade is 1, hence CI1.] The element
abundance curve for the solar photosphere, for at least the
dozen most abundant elements, is the same as that for other
stars in the galaxy and nearby galaxies. Thus, the principles
developed below are applicable for exoplanet building too.
As building blocks, chondrites provide the basis for com-
positional models that can plausibly describe what is in the
Earths core. Some proposed models of the Earths composi-
tion fall outside of the observed range for chondrites, with
enrichments in refractory elements and stronger depletions in
volatile element abundances than seen in chondrites (Fig. 1)
(McDonough 2014). Element classication (i.e., volatile ver-
sus refractory) is dened in terms of half-mass condensation
temperatures of an element during cooling of a nebular accre-
tion disk and precipitation into common mineral phases.
Given the plausible range of volatile and refractory element
abundances in different groups of chondrites that accreted to
Earths Core 3
form the Earth, it is not possible to uniquely dene the
planets composition. Moreover, while chondrites represent
a compositional guide, the Earth is unlikely to match the
composition of any specic, single chondrite. The standard
rules for planet building are as follows:
Use orbital and seismic constraints, if available, to dene
the physical state
Use chemical and isotopic data from chondritic (i.e., prim-
itive) meteorites and the solar photosphere
Assume chondritic proportions for the refractory elements
and model their absolute abundances based on chemical
trends observed for mantle samples (basalts and
peridotites),
Model the absolute abundances of Mg, Fe, Si, and O; these
are the major elements, which are nonrefractory elements,
and constitute >90% by mass of the rocky planets
Recognize a chemical gradient in solar system (e.g., rocky
planets versus the gas giants)
Model the relative and absolute abundances of the non-
refractory elements in the planet
Model core compositions are estimated by subtracting the
composition of the silicate Earth from the composition of the
bulk planet. The relative distribution of lithophile elements
(those that tend to bond with oxygen and are excluded from
the core) versus siderophile elements (those that tend to bond
with iron and are dominantly stored in the core) in the crust
and mantle dene the composition of the silicate Earth. For
refractory elements, including lithophile and siderophile,
their relative proportions in the planet are the same as in
chondrites (i.e., the lithophiles have constant ratios at !10%
or better, constant ratios of siderophiles are at !15% or better,
and ratios of lithophile to siderophile have uncertainties at
circa !25%, with the greater difference for the latter being
due to variations in the proportion of metal to silicate in
chondrites). Therefore, if we can predict the abundance of
one refractory element in the Earth, we can predict the plan-
etary inventory of all 37 refractory elements, which include:
Be, Al, Ca, Sc, Ti, V, Sr, Y, Zr, Nb, Ba, REE, Hf, Ta, Th, and
U (these are the lithophiles); and Mo, Ru, Rh, W, Re, Os, Ir,
and Pt (these are the siderophiles, with V showing both
behavior). How and why elements are distributed between
the core (mostly siderophile) and the surrounding silicate
shell (lithophile) depends on the conditions of pressure, tem-
perature, and redox state during core formation. The condi-
tions of core formation are critical to establishing its nal
composition.
Among the dozen most common elements in the solar
system, the rocky planets are variably depleted in the most
volatile constituents (i.e., the gases: H, He, C, N, O, Ne, Ar).
The remaining common elements are O, Fe, Si, and Mg,
which constitute between 90% and 95% the mass of the
Earth, Mars, and Venus roughly in the mass proportion
30:30:20:20 (or in the approximate atomic proportions of
50:15:15:15), respectively (Fig. 2). The high density and
small radius of Mercury indicates that it likely has a much
higher proportion of iron than the other rocky planets. Com-
positional models for the Earth, other terrestrial planets, and
exoplanets with bulk density in the range 30006000 kg/m
3
will be dictated by mixtures of metal and silicate. The silicates
are predominately a bimineralic mix of olivine and pyroxene
at pressures equivalent to the Earths upper mantle. At the
seismically dene base of the Transition Zone (i.e., 660 km
depth, ~22 GPa), olivine disproportionates into bridgmanite
(Mg-Fe silicate perovskite) and ferro-periclase, whereas
the pyroxene component (which in the upper mantle has
gradually dissolved into a majoritic garnet) converts to
bridgmanite and Ca perovskite over a broader range of pres-
sures (i.e, ~1924 GPa).
A big unknown in predicting the bulk composition of a
planet is establishing the proportion of olivine to pyroxene
Earths Core, Fig. 1 Chondrite normalize compositions for carbona-
ceous chondrites (CM, CO, CV subtypes), ordinary chondrites (H, L,
and LL subgroups), enstatite chondrites (EH and EL subgroups) and the
Earth. The gure is double normalized with all compositional models
plotted as having the same Si content, which corrects for differing
amounts of highly volatile element contents (i.e., H, C, N, O). Normal-
izing reference frame assumes a CI1 chondrites composition. Data
for the chondrites are from Wasson and Kallemeyn (1988) and
McDonough (2014)
4 Earths Core
accreted by the body, which is indicated by its Mg/Si ratio.
There is no xed Mg/Si ratio for chondrites and thus there is
no chondritic constraint (Fig. 3). Olivine has a 2:1 molar
Mg/Si content, whereas pyroxenes molar ratio is 1:1. Astro-
mineralogical observations of accretion disks reveal radial
variations in the proportion of olivine to pyroxene, some
with inner disk regions being richer in olivine relative to the
disk wide composition (van Boekel et al. 2004), whereas
other have more olivine in the outer part of the circumstellar
disk; differences in disk mineralogy may relate to type of star
(e.g., T Tauri vs. Herbig Ae/Be stars) (Bouwman et al. 2010).
Likewise, one can also envisage disk evolution in time with
temperature decreasing as the disk cools leading to decreasing
Mg/Si (less olivine ppt with time). Moreover, crystallinity
scales radially, with inner disk regions (a few AU) having
higher abundances of large grains and generally higher crys-
tallinity as compared to outer disk regions, suggesting grain
growth occurs more rapidly in the inner disk regions
(DAlessio et al. 2005; Sargent et al. 2009). Thus, planetary
modelers must acknowledge that there is no xed Mg/Si value
for a planet.
After O, Fe, Si, and Mg, the next most abundant elements
in chondrites and rocky planets are Al, Ca, and Ni and
together these seven elements, plus S, make up >98% by
mass of these solid bodies. Both Ca and Al are refractory
elements, and their planetary weight ratio is 1.1 (atomic ratio
is 0.72), equal to that of chondrites, but their absolute abun-
dances in the Earth are model dependent; these elements are
exclusively lithophile and thus concentrate in only the silicate
portion of the Earth. Nickel and sulfur are both nonrefractory
elements and are partitioned into both the core and mantle,
with Ni being dominantly siderophile (>90% of the Earths
Ni is hosted in the core) and sulfur is the denition of what is a
chalcophile element (>95% of the Earths sulfur is in the
core) (McDonough 2014).
The ratio of metals to oxides is a basic feature used to
classify chondrites. The chondrites, carbonaceous (highly
oxidized), ordinary (an intermediate redox state), and
enstatite (highly reduced state) can be divided into subgroups
Earths Core, Fig. 2 Atomic
proportions of O, Fe, Si, and Mg in
chondrites and the Earth. Data
sources listed in Fig. 1. The
O context is #50% in the
chondrites and the total atomic
fraction for these four elements in
all of these materials adds up to
>90%
Earths Core, Fig. 3 The compositional space for the chondrites and
the Earth in a tetrahedral volume plotted from the apices of Si, Mg, Fe,
and O, with the data plotted onto the surface of the Si-Fe-Mg plane. Data
sources listed in Fig. 1. The Earths composition falls outside of the eld
of the known chondrites. This model composition for the Earth is
acceptable given observations of the spatial variations in rates of olivine
to pyroxene crystallization in accretion disks (See text for further details)
Earths Core 5
of high, low, and low-low iron contents. Therefore, the
amount of Fe relative to O, Si, and Mg in chondrites and
planets can vary considerably. The redox state of a planet can
be readily established by determining the Mg # (atomic
Mg/(Mg + Fe)) of mantle silicates in combination with a
value for the mass proportion of metallic to silicate shells.
The Earth with 90% of its iron in metallic state is a reduced
body with a redox state intermediate between ordinary and
enstatite chondrites (Fig. 4).
Importantly, there are a number of critical element ratios
for the nonrefractory elements that are relatively constant in
chondrites and, if these element pairs have similar condensa-
tion temperatures, then these xed ratios can be used to place
compositional constraints on the bulk planet. The xed chon-
dritic ratios of Fe/Ni (17.4 !0.5) and Ni/Co (20.5 !0.6)
constrains the planetary and core abundances of Ni and Co,
given well established values for the mantle. To this end, a
and Li and Agee (1996) identied the pressure dependence
partitioning of Ni relative to Co and found that at mid-mantle
conditions the relative partitioning of these two elements
approached 1, with both the core and mantle having chon-
dritic Ni/Co values. This nding was fundamental as the
silicate Earth and core are assumed to have a mass balance
in these ratios.
The initial growth of Earth and core-mantle segregation is
often consider in terms of occurring under conditions of
homogeneous or heterogeneous accretion, with the former
model assuming that the accreting materials had a xed,
constant composition throughout planetary growth, while
latter model models envisages signicant changes in the
composition of material that accretes to the Earth as it grew
models seek to explain the siderophile element composition
of the mantle. For the most part models of homogeneous
accretion require a continuous and restricted set of physical
and chemical conditions to achieve the present composition
of the core and mantle, whereas heterogeneous accretion
models allow for a range of probable conditions envisaged
in an evolving proto-planetary disk, including temporal var-
iations in redox conditions and the relative proportions of
volatile to refractory materials.
Given an average condition for temperature, pressure, and
oxygen fugacity of core formation we can construct models
for the core that t the existing constraints and match com-
positional models of the silicate Earth and chondrites.
The abundance, absolute and relative, of various redox
sensitive siderophile elements (e.g., W, V, Cr, Nb, Mn, Mo,
and others) in the silicate Earth further dene the relative
silicate-metal partitioning of elements during core formation
(Badro et al. 2015; Badro et al. 2014; Corgne et al. 2008;
Fischer et al. 2015; Rubie et al. 2015). At this stage there
remains one further consideration before estimating the
Earths Core, Fig. 4 A plot of the oxidized iron to Si content of
chondrites versus the metallic and sulfur-bearing content to Si content
of chondrites. This diagram is also referred to as a Urey-Craig diagram.
The Earth plots intermediate to the highly reduced enstatite chondrites
and the less reduced ordinary chondrites. The CM and CI chondrites
have nearly (or all) iron as oxidized, with no metallic iron
6 Earths Core
composition of the core, and that is identifying the chemical
identity of the element(s) accounting for its density decit.
Finally, an important constraint on core formation came
from the observation that the primitive mantle abundances of
the highly siderophile elements (i.e., those with metal/silicate
partition coefcients of 10
4
and greater and include the plat-
inum group metals (Ru, Rh, Pd, Os, Ir, Pt), Re and Au) are in
chondritic proportions and are enriched in the mantle at con-
centration levels well beyond that expected for their known
metal/silicate partition coefcients (Kimura et al. 1974). Fol-
lowing a range of experimental studies it has been demon-
strated that that the relative metal/silicate partitioning
behavior of the highly siderophile elements are 10
4
and
greater and that the relative differences between different
member of this element group would result in grossly non-
chondritic ratios for the Re-Os-Pt isotope system, which is not
observed (Brenan et al. 2016; Day et al. 2016). Thus, the
mantle signature requires the addition of a late veneer of
added chondritic material, with a mass contribution equiva-
lent to about 0.5% to <1% the mass of the mantle (Day
et al. 2016).
Core Composition and Its Light Element(s)
Table 2 presents a model for the composition of the Earth, its
core, and the silicate Earth. The core model was constructed
as the difference based on estimates of the bulk silicate Earth
and the bulk planet compositions. The core composition
proposed here has a mixture of O, Si, and S that accounts
for the density decit (37%). The relative proportions of
O and Si in the core were established by metal/silicate
partitioning (Badro et al. 2015; Fischer et al. 2015; Rubie
et al. 2011). Assuming the cores mean atomic number to be
23 (Birch 1964), the absolute proportions of O and Si is then
xed. By way of comparison, McDonough (2014) proposed
two end-member models, with one model containing Si and
S and no O as the light elements and the other model contains
O and S and no Si. Other compositional models are presented
for comparison in Table 1in Hirose et al. (2013), where they
have sorted the different models in terms of simple single light
element versus multiply, mixed light element models. One
should view all of these models as acceptable, given the
available constraints.
Let me start by saying that I do not know what is (are) the
light element(s) in the core. The issue has dogged our science
for half of a century and will continue to do so for some time.
Words of caution are needed before discussing further the
composition of the core. My bias as a geochemist states that
observations on the composition of the silicate Earth and the
range of compositions in chondrites denes the recipe that
must be matched by experimental studies to dene the P-T-
fO
2
-Xi (i.e., pressure, temperature, oxygen fugacity, and
compositional) space for metal-silicate equilibrium during
core formation. Models describing the dynamics of core for-
mation and its nal composition are probability positions and
need to be recognized as such. As a second cautionary note,
the periodic table is not a one-stop shopping table, if it is
argued that a minor or trace element is in the core, then other
elements in the same column are likely there too in some
proportional relationship (e.g., Cr: Mo and W, Mn: Re, V: Nb
and Ta). This principle is fundamental chemistry reecting the
importance of outer orbital electrons in controlling the
partitioning behavior of elements. Thirdly, all partition coef-
cients (D-values) go to 1 as temperature goes to innity and
thus metal-silicate fractionation becomes muted at higher
temperature conditions. Finally, appeals to adding specic
elements into the core are also understandably done so to
explain the power driving the geodynamo; resolving the
energy equation for the geodynamo is intractable and will
be so for some time.
The identity of the light element(s) in the core is elusive.
Many suggestions have been published over the decades
(e.g., H, C, N, O, Mg, Si, S). In general, any element with a
lower atomic number than
26
Fe is a likely candidate. Assum-
ing Birchs law
Vf¼K
r
!"
0:5
;Vp¼MrþBþAT&T0
ðÞr&r)
ðÞ(1)
where Vfis hydrodynamic or bulk sound velocity, V
P
is
compressional wave velocity, kis bulk modulus, ris density,
Mand Bare room temperature coefcients of Birchs law, and
Sakamaki et al. (2016))], one can predict the mean atomic
weight of the mantle or core. Recently, Helffrich (2015)
reported a similar relationship between sound speed and
molecular weight using a hard sphere model for liquids. In
addition, Poirier (1994), Anderson and Isaak (2002) and
Badro et al. (2014) present simple graphical models to deter-
mine the amount (in wt%) of light element mixture in the core
that scales as a function of core temperature. The model
presented in Table 2uses these principles to constrain the
proportion of O and Si in the core. The S content of the core is
established on cosmochemical grounds (Dreibus and Palme
1996) and constraints from the modeling of the bulk silicate
Earth (McDonough and Sun 1995).
Chemical and physical observations are useful for devel-
oping plausible models for the nature and amount of a light
element(s) in the core. Lithophile elements like O, Si, and Mg,
in addition to Fe, are the most abundant elements in the Earth
and can be readily defended as being dissolved in a core
forming metallic liquid. Likewise, the gases of H-C-N-O are
abundant in the solar system and would have been abundant
in the nebular disk. The rocky planets might have had a dense
Earths Core 7
surrounding atmosphere, particularly early in their evolution,
which increases the solubility of these gases in a magma
ocean and the potential for partitioning these elements into a
core forming metallic melt. Thus, models invoking H, C,
and/or N as constituents of the core cannot be dismissed. On
the other hand, however, the Earth, and potentially the rest of
the terrestrial planets in the solar system, is volatile depleted,
as documented by its Rb-Sr, K-Ar, K-U, and He-Ne-Ar iso-
topic systems and total radiogenic heat production. Therefore,
the Earths volatile depletion curve establishes the limit of 2%
(by mass) of sulfur in the core. Overall, chemistry offers
guidelines, but a few constraints in eliminating competing
core compositional models. Physical observations provide
the best limits on acceptable core models.
The small addition of the lightest elements (e.g., H and C)
is appealing because they offer considerable leverage in low-
ering the mean atomic weight of the core. It is noteworthy that
the core composition reported in Table 2posits signicant
amounts of C and H in the core given the Earths volatility
curve (McDonough 2014). However, their appeal is dimin-
ished if their addition to the core also comes with conse-
quences for the mantle (e.g., limited metal-silicate
fractionation for H and C during core formation). Composi-
tional models for the mantle have proposed a range of H and
C abundances (Dasgupta and Hirschmann 2010; Halliday
2013; Marty 2012) and variable estimates for the contents of
H and C in the core and some model compositions proposed
for the Earths core suggest abundant carbon (Wood 1993) or
hydrogen (Williams and Hemley 2001) for consideration.
Considerable attention has been paid to models with O, Si,
or both as likely light element candidates (Badro et al. 2015;
Fischer et al. 2015; Li and Fei 2014; Ringwood 1984). Based
on the differences in the silicon isotopic composition of
chondrites and the Earth, Fitoussi and Bourdon (2012) and
Fitoussi et al. (2016) proposed that the Earth core contains
some 47 wt% Si, with the range taking into account metal-
silicate isotopic fractionation at presumed core separation
temperatures. Initially, the presence of silicon and oxygen in
the core was considered an either or condition, whereas today
many have found a range of acceptable conditions of core
formation that result in iron alloys with mixed amounts of
silicon and oxygen. More recent models suggest that Mg may
have once been present as a light element in the core (Badro
et al. 2016;ORourke and Stevenson 2016). This suggestion
is permitted given that Mg is one of the four most abundant
elements in the Earth. Moreover, these authors note that as the
core cools down it exsolves MgO (or Mg silicate) and this
reaction is exothermic and thus provides a power source for
driving the geodynamo. Alternatively, Hirose et al. (2017)
have suggested that exsolution of SiO
2
from the outer core is
another potential power source for the geodynamo. The
present-day situation we face is that there are many competing
models to explain the core composition and potentially also
energy sources for driving the cores dynamo.
Power to Drive the Geodynamo and Radioactive
Elements in the Core
The geodynamo is produced in the outer core and demands
power for its operation. Several papers have considered the
presence of radioactive, heat producing elements (i.e.,
Earths Core, Table 2 Composition of the Earth, silicate Earth, and core in weight percent and atomic proportions
wt% Earth Silicate Earth Core Atomic prop. Earth Silicate Earth Core
Fe 31.9 6.26 85.5 Fe 0.148 0.024 0.750
O 30.4 44 2 O 0.493 0.581 0.061
Si 15.3 21 4 Si 0.142 0.158 0.070
Mg 15.4 22.8 0 Mg 0.164 0.198 0
Ni 1.83 0.196 5.10 Ni 0.008 0.001 0.043
Co 0.090 0.0105 0.249 Co 0.0004 0.00004 0.002
Ca 1.71 2.53 0 Ca 0.011 0.013 0
Al 1.59 2.35 0 Al 0.015 0.018 0
S 0.650 0.025 1.9 S 0.005 0.0002 0.029
Cr 0.43 0.263 0.75 Cr 0.002 0.001 0.007
Na 0.18 0.27 0 Na 0.002 0.002 0
P 0.072 0.009 0.02 P 0.001 0.0001 0.0003
Mn 0.08 0.105 0.03 Mn 0.0004 0.0004 0.0003
C 0.073 0.012 0.2 C 0.002 0.0002 0.008
H 0.026 0.01 0.06 H 0.007 0.002 0.029
Total 99.7 99.8 99.8 Total 1.000 1.000 1.000
Absolute and relative proportions of O and Si in the core were established by metal/silicate partitioning and xing the cores mean atomic number to
22.9 (c.f. Birch 1964)
8 Earths Core
HPE =K, Th, and U) in the Earths core, as they provide
power to drive the cores dynamo. There is no compelling
experimental or geochemical constraint that requires HPE in
the core. Moreover, doing so often requires incorporation of
other lithophile elements for which there is no supporting
evidence for their depletion in the silicate Earth.
The energy budget of the core does not require HPE, as
alternative energy sources (e.g., oxide exsolution, see above)
may also provide the needed power. In addition, unknowns in
understanding the generation and maintenance of a
geodynamo are great, which includes being able to describe
accurately the cores energy equation.
As is the case with the discussion of the light element in the
core, so too for the presence of radioactive elements in the
core: there are big unknowns. Thus, one should not confuse
the conditions of what is plausible and what is needed.
Potassium: Incorporation of K in the core has been con-
sidered by numerous workers (Corgne et al. 2007; Li and Fei
2014; Watanabe et al. 2014) and generally found that a few to
a couple tens of mg/g of K, or less, can be accommodated in
the core. Using ab initio calculations Lee et al. (2004)
suggested that the metallization of potassium at pressure,
due to s-to d-orbital transitions, as a mechanism for alloying
potassium with Fe. Both Cs and Rb, larger alkali metals,
would undergo similar orbital transitions, but at lower pres-
sures. There is no evidence, however, that Rb is anomalously
depleted in the mantle with respect to the planetary volatility
curve, contrary to an expectation from the pressure-dependent
metallization process.
Uranium and thorium: Wheeler et al. (2006) and
Malavergne et al. (2007) show that under core segregation
conditions U remains strongly lithophile and is for the
most part excluded from the core. Similarly, thorium has not
been found to partition into a core forming metal phase; it
is understood to be wholly lithophile. An additional constraint
on the abundance and relative distribution of Th and U in
the bulk silicate Earth comes from Pb isotope studies. The
k
Pb
ratio (i.e., the time integrated Th/U ratio derived from the
208
)
Pb/
206
)
Pb ratio) of the silicate Earth is comparable to that
of chondrites at !25% (Andersen et al. 2015; Castillo 2016;
Elliott et al. 1999; Galer and ONions 1985; Kumari et al.
2016; Paul et al. 2003), consistent with negligible amounts of
Th and U in the Earths core. Studies on the partitioning of
thorium between metal and silicate demonstrate that it is
wholly lithophile in its behavior.
The Age of Core-Mantle Separation
Age of core formation is constrained by the
182
Hf-
182
W
isotope system (t
1/2
of 9 Ma) and known to be on the order
of 30150 million years after the age of solar system forma-
tion (Kleine et al. 2002; Kleine et al. 2009; Yin et al. 2002).
The solar system formation age (t
0
=4568 Ga) is established
mostly from the absolute system, U-Pb isotopes, and the
relative system
26
Al/
26
Mg isotopes (t
1/2
of 0.7 Ma), as applied
to the earliest condensates, which are high temperature,
refractory materials (i.e., CAIs, calcium alumina inclusions)
found as inclusions in chondritic meteorites (Brennecka et al.
2015; Connelly et al. 2012; MacPherson et al. 2012).
Together these isotopic systems indicate that the bulk of
Earths accretion occurred some 10
6
10
7
years after t
0
and
core formation followed as accretion progressed, including
some modication during and following the Moon-forming
impact event.
The current core formation paradigm posits segregation of
sinking metal diapirs that chemically interact with the imme-
diately surrounding silicate (liquid or solid) during or follow-
ing a magma ocean(s) environment of the early Earth. The
degree of metal-silicate equilibrium occurring during descent
of these diapirs is unknown. However, as metal aggregates
ponds and then begin to sink the potential for chemical
exchange becomes less efcient as the length scale increases.
Chemical consequences of a large Moon-forming impact event
introducesother unknowns, for which we donot know whether
or not the impactors core merged, emulsied in the mantle, or
otherwise superimposed additional complications on the cores
composition (Nakajima and Stevenson 2015).
Hafnium and tungsten readily and effectively separate
between the metallic (core) and silicate Earth (planet minus
the core), with Hf hosted nearly exclusively in the silicate
Earth and W almost wholly (90%) in the core. Due to the solar
system initiating supernova, the solar system (and hence the
Earth) inherited a fraction of live
182
Hf (Lugaro et al. 2014)
that decayed into
182
W and consequently the age of core
formation reects the competing processes of metal segrega-
tion versus
182
Hf decay.
Chronologically, what has been established recently is that
small bodies, as represented by groups of iron meteorites,
accreted and differentiated quickly (i.e., 05 Ma after t
0
)
(Kruijer et al. 2014). During the waning stages of small-
body accretion (t
0
+35 million years), and when radiogenic
heating from the
26
Al/
26
Mg isotope system was running out
of energy, the formation of chondritic bodies occurs, which
are the most primitive and undifferentiated planetesimals in
the solar system. Hence, the preservation of chondritic parent
bodies represents a waning stage of accretion of small bodies.
It is likely that Earths formation occurred throughout this
sequence and beyond and consequently its inventory of
182
W
was shared between the core and mantle, revealing that the
amount of
182
W in the silicate Earth establishes a lower limit
to core formation age, while the timing of Moon formation
sets the upper limit. Our present estimate is the silicate Earth
has a m
182
W isotopic composition of zero (it is the reference
point for the isotope system and units are expressed as devi-
ations in parts in 10
6
), chondrites are approximately &200
Earths Core 9
and the Earths core is slightly more negative (~&220)
(Walker 2016).
Inner Core and Its Age of Crystallization
The inner core is <2% the mass of the Earth and 5% the mass
of the core. It is slightly smaller than the Earths moon and is
at about the temperature of the Suns surface. The ICB (inner-
outer core boundary) surface provides a constraint for esti-
mating the temperature at the center of the Earth. An upper
temperature limit is obtained from the melting temperature of
Fe at relevant temperatures and pressures (~6200 K, 330 GPa)
and an assumed conductive thermal gradient (Fischer 2016).
Experimental challenges associated with determining the
melting curve of iron and the upper temperature limit of the
ICB are formidable; there is a range of 1500 K at 330 GPa for
the Fe-alloy melting curves (Fischer 2016). Nonetheless,
given the potential light element candidates and their inu-
ence on melting point depression, an ICB temperature of
5700 !550 K covers the range of estimates (Fischer 2016).
The inner core has about half the density decit relative to
that of the outer core (Anderson and Isaak 2002; Jephcoat and
Olson 1987; Masters and Gubbins 2003); proportionally, one
might intuit that this translates to approximately half the
amount of light element. The phase diagram dictates that the
inner core crystallization temperature will be lower than that
for pure iron due to freezing point depression. Thus, an
estimate of the temperature of the ICB depends on ones
estimate for the light element composition of the outer core,
which is a concatenation of unknowns.
Compositional estimates for the inner core are not forth-
coming, as the unknowns are great, and come with consider-
able nonuniqueness (Antonangeli et al. 2010; Badro et al.
2007; Sakamaki et al. 2016). We know that its solid, likely
in an hcp (hexagonal close packed) structure, there is
820 !180 kg/m
3
density increase across the ICB (Masters
and Gubbins, 2003), it has a density decit, and it is aniso-
tropic (polar vs equatorial wave speed). The inner core may be
undergoing thermal convection (Cottaar and Buffett 2012),
which may explain its seismic anisotropy.
Solidication of the inner core produces latent heat and
results in a buoyant, residual liquid enriched in light elements.
The Coriolis force from the Earths rotation, combined with a
heated, compositionally buoyant uid, gives this ascending
uid the power to generate a magnetic eld. In addition, heat
extracted from the core by the mantle, as the planet cools,
gives rise to the cores thermal convection. In combination
with rotational forces, the thermal and compositional effects
are the most probable energy sources driving the geodynamo.
When did inner core crystallization begin? This question
hinges greatly on understanding the thermal history of the
core. There are many unknowns that contribute to this
question and most are active areas of research accompanied
by considerable debate, these include: the thermal and elec-
trical conductivity of core materials (where there is an ongo-
ing debate over the factor of 2 difference in the assumed
values for these parameters), the initial and present day core
temperature, heat ow across the CMB (see Table 1), and
whether there are (and how much) radioactive elements in the
core. Debates on these topics will continue for some time as
technical and computational challenges are great, including
understanding the energy budget needed to drive the
geodynamo.
inner core crystallization, as early as 4.5 Ga. Models of early
inner core growth were nearly exclusively based on anoma-
lous
186
Os-
188
Os and
187
Os-
188
Os isotopic compositions in
basalts whose source region experience core additions; the
origins of these anomalies came from decays in the Re-Os and
Pt-Os isotope systems and parent daughter fractionation due
to inner core crystallization (Walker 2016). More recently,
however, corroborating
182
W isotopic evidence of this core-
basalts and thus the model of early inner core growth based on
isotopic evidence has been abandoned (Walker 2016).
Geophysical estimates for the age of the inner core are
based on a full understanding of the cores energy equation
and usually includes a trade-off between internal heat produc-
tion and heat dissipation across the core-mantle boundary.
Labrosse et al. (2001) examined the question assuming a
role for heat generation from potassium, modeling up to
200 mg/g K in the core, and showed that core crystallization
might have started anytime from the Archean to nearly the
present depending on how much potassium is in the core.
Later, addressing this issue following the recently revised
thermal and electrical conductivity values for the core,
Gomi et al. (2013) and Labrosse (2015) concluded that the
amount of potassium in the core plays less of a role in
establishing the initiation of inner-core crystallization, but
that the core cooling rate is signicant. From these analyses
a young age (circa ~05 Ga) for initiation of inner core is
favored. Alternatively, estimates of an inner core having
begun crystallization from 1.5 to 0.5 billion years ago are
common (Labrosse et al. 2001). Overall, there are few con-
straints regarding the physical state and evolution of the core
and inner core system, leaving us with open questions about
when the inner core began to crystallize, the heat ux across
the core-mantle boundary, and the power budget for the
geodynamo.
Summary
The Earths core comprises the innermost 3483 km of the
planet. Both the inner solid core and outer liquid core consist
10 Earths Core
of about 85% iron and 5% nickel with *10% other elements
including lighter elements O, Si, and/or S as well as most of
the inventory of highly siderophile elements such as the
platinum group elements and gold. A metallic liquid core
had likely formed within the rst 3060 million years of
solar system history. The solid inner core has grown by
crystallization of the outer core, a process that may have
begun as recently as 0.51.5 billion years ago.
Cross-References
Chondrites
Cosmic Element Abundances
Earths Formation and Evolution
Giant Impact Hypothesis
Heat Producing Elements
Iron
Iron Meteorites
Lithophile Elements
Moon Formation
Nickel
Oxygen Fugacity
Partition Coefcients
Siderophile Elements
Sulfur
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... This may inevitably underestimate the core mass and thus overestimate core radius should the planet's core be differentiated. However, we envisage that the under-/over-estimation should not be significant, concerning that Earth's inner solid core just accounts for 5% of the mass of the core (Yoder 1995;McDonough 2017) and that even a planet as small as Mars has been suggested to be in a molten state based on the most recent seismic data from the InSight mission (Stähler et al. 2021). Conflicting views on liquid/solid cores for super-Earths have been presented (e.g. ...
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A major goal in the discovery and characterisation of exoplanets is to identify terrestrial-type worlds that are similar to (or otherwise distinct from) our Earth. Recent results have highlighted the importance of applying devolatilisation -- i.e. depletion of volatiles -- to the chemical composition of planet-hosting stars to constrain bulk composition and interiors of terrestrial-type exoplanets. In this work, we apply such an approach to a selected sample of 13 planet-hosting Sun-like stars, for which high-precision photospheric abundances have been determined in the first paper of the series. With the resultant devolatilised stellar composition (i.e. the model planetary bulk composition) as well as other constraints including mass and radius, we model the detailed mineralogy and interior structure of hypothetical, habitable-zone terrestrial planets ("exo-Earths") around these stars. Model output shows that most of these exo-Earths are expected to have broadly Earth-like composition and interior structure, consistent with conclusions derived independently from analysis of polluted white dwarfs. The exceptions are the Kepler-10 and Kepler-37 exo-Earths, which we predict are strongly oxidised and thus would develop metallic cores much smaller than Earth. Investigating our devolatilisation model at its extremes as well as varying planetary mass and radius (within the terrestrial regime) reveals potential diversities in the interiors of terrestrial planets. By considering (i) high-precision stellar abundances, (ii) devolatilisation, and (iii) planetary mass and radius holistically, this work represents essential steps to explore the detailed mineralogy and interior structure of terrestrial-type exoplanets, which in turn are fundamental for our understanding of planetary dynamics and long-term evolution.
... These same studies either did not include Th in their experiments or report DU partition coefficients and below limits of detection values for DTh (Chidester et al., 2017). The Earth's core is recognized as having a mix of light elements that account for its density deficit relative to iron, with a limited contribution from the volatile element sulfur (McDonough, 2017). In a sulfur saturated, non-peridotitic set of experiments, DTh was found to be ~0.1DU ...
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Radioactive decay of potassium (K), thorium (Th), and uranium (U) power the Earth's engine, with variations in 232Th/238U recording planetary differentiation, atmospheric oxidation, and biologically mediated processes. We report several thousand $^{232}$Th/$^{238}$U ($\kappa$) and time-integrated Pb isotopic ($\kappa$$_{Pb}) values and assess their ratios for the Earth, core, and silicate Earth. Complementary bulk silicate Earth domains (i.e., continental crust \kappa_{Pb}^{CC} = 3.94 ^{+0.20}_{-0.11} and modern mantle \kappa_{Pb}^{MM} = 3.87 ^{+0.15}_{-0.07}, respectively) tightly bracket the solar system initial \kappa_{Pb}^{SS} = 3.890 \pm 0.015. These findings reveal the bulk silicate Earth's \kappa$$_{Pb}^{BSE}$ is 3.90 $^{+0.13}_{-0.07}$ (or Th/U = 3.77 for the mass ratio), which resolves a long-standing debate regarding the Earth's Th/U value. We performed a Monte Carlo simulation to calculate the $\kappa_{Pb}$ of the BSE and bulk Earth for a range of U concentrations in the core (from 0 to 10 ng/g). Comparison of our results with $\kappa$$_{Pb}^{SS}$ constrains the available U and Th budget in the core. Negligible Th/U fractionation accompanied accretion, core formation, and crust - mantle differentiation, and trivial amounts of these elements (0.07 ppb by weight, equivalent to 0.014 TW of radiogenic power) were added to the core and do not power the geodynamo.
Thesis
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Comparative studies of terrestrial planets have provided important insights into physico-chemical processes that produced their similarities and differences. Chemical composition of terrestrial planets records planetary accretion, differentiation, impact, and surface processes. In order to place new constraints on the origin and evolution of rocky planets, I investigated chemical compositions of terrestrial planets and meteorites from primitive asteroids (chondrites). Compositional models of planets have found or assumed chondritic relative abundances of refractory lithophile elements (RLE). I challenge this fundamental paradigm by showing highly variable RLE ratios in individual chondrules from enstatite chondrites (EC), which are highly reduced primitive meteorites with Earth-like isotopic signatures. The fractionated RLE compo- sitions of EC chondrules reflect moderately chalcophile behaviors of these elements and sulfide-silicate separation in highly reduced nebular environments. If the Earth’s building blocks were dominated by highly reduced EC-like materials, they should not have been affected by a physical sorting of silicates and sulfides before their accretion. Alternatively, the Earth’s precursors might have been high-temperature nebular materials that condensed before precipitation of the RLE-bearing sulfides. The bulk Mercury might not have chondritic RLE ratios, due to sulfide-silicate separation processes that formed its large metallic core. Previous compositional models of Mars relied on an assumption of CI-chondritic relative abundance of Mn and more refractory elements, which has been challenged by recent astrophysical observations. Here I propose a new martian model composition that avoids such an assumption, using data from martian meteorites and spacecraft observations. The new model finds that Mars is enriched in refractory elements and show a systematic depletion of moderately volatiles as a function of their volatilities compared to the CI abundance. The Mars’ volatile depletion trend indicates a S-poor composition for the martian core, which requires an incorporation of additional light elements (e.g., O, H) into the core to match the martian geodetic properties. Earth and Mars are equally enriched in refractory elements, although Earth is more volatile-depleted and less oxidized than Mars. These compositional properties were established by a nebular fractionation, with negligible post-accretionary losses of moderately volatile elements. The degree of planetary and asteroidal volatile element depletion might correlate with the abundances of chondrules in the accreted materials, planetary size, and their accretion timescale. During its prolonged formation, the Earth likely accreted more chondrules and less matrix-like materials than Mars and chondritic asteroids. The correlations between these planetary properties constrain the composition and origin of Mercury, Venus, the Moon-forming giant impactor, and the proto-Earth. These observations, combined with the shared refractory enrichment in Earth and Mars, and insights from planetary uncompressed densities, establishes the compositional model of Mercury. Uncompressed den- sities of rocky bodies in the solar system decrease with their heliocentric distance, indicating a role of disk-scale metal-silicate separation before the planetary accretion, rather than post- accretionary modification processes. Insights presented here update our knowledge of the origin of rocky planets in our solar system, and provide future perspectives on studies of chemistry of (extra)solar system bodies.
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Debate continues on the amount and distribution of radioactive heat producing elements (i.e., U, Th, and K) in the Earth, with estimates for mantle heat production varying by an order of magnitude. Constraints on the bulk‐silicate Earth's (BSE) radiogenic power also places constraints on overall BSE composition. Geoneutrino detection is a direct measure of the Earth's decay rate of Th and U. The geoneutrino signal has contributions from the local ( ∼ 40%) and global ( ∼ 35%) continental lithosphere and the underlying inaccessible mantle ( ∼ 25%). Geophysical models are combined with geochemical data sets to predict the geoneutrino signal at current and future geoneutrino detectors. We propagated uncertainties, both chemical and physical, through Monte Carlo methods. Estimated total signal uncertainties are on the order of ∼ 20%, proportionally with geophysical and geochemical inputs contributing ∼ 30% and ∼ 70%, respectively. We find that estimated signals, calculated using CRUST2.0, CRUST1.0, and LITHO1.0, are within physical uncertainty of each other, suggesting that the choice of underlying geophysical model will not change results significantly, but will shift the central value by up to ∼ 15%. Similarly, we see no significant difference between calculated layer abundances and bulk crustal heat production when using these geophysical models. The bulk crustal heat production is calculated as 7 ± 2 TW, which includes an increase of 1 TW in uncertainty relative to previous studies. Combination of our predicted lithospheric signal with measured signals yield an estimated BSE heat production of 21.5 ± 10.4 TW. Future improvements, including uncertainty attribution and near‐field modeling, are discussed.
Preprint
Exoplanet interior modelling usually makes the assumption that the elemental abundances of a planet are identical to those of its host star. Host stellar abundances are good proxies of planetary abundances, but only for refractory elements. This is particularly true for terrestrial planets, as evidenced by the relative differences in bulk chemical composition between the Sun and the Earth and other inner solar system bodies. The elemental abundances of a planet host star must therefore be devolatilised in order to correctly represent the bulk chemical composition of its terrestrial planets. Furthermore, nickel and light elements make an important contribution alongside iron to the core of terrestrial planets. We therefore adopt an extended chemical network of the core, constrained by an Fe/Ni ratio of 18 $\pm$ 4 (by number). By applying these constraints to the Sun, our modelling reproduces the composition of the mantle and core, as well as the core mass fraction of the Earth. We also apply our modelling to four exoplanet host stars with precisely measured elemental abundances: Kepler-10, Kepler-20, Kepler-21 and Kepler-100. If these stars would also host terrestrial planets in their habitable zone, we find that such planets orbiting Kepler-21 would be the most Earth-like, while those orbiting Kepler-10 would be the least. To assess the similarity of a rocky exoplanet to the Earth in terms of interior composition and structure, high-precision host stellar abundances are critical. Our modelling implies that abundance uncertainties should be better than $\sim$ 0.04 dex for such an assessment to be made.
Chapter
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This chapter examines the composition of the core and of the mantle and its domains, upper and lower, its physical and chemical attributes, and its evolution. It starts with fundamental definitions, particularly of what is the lower and upper mantle. Although a compositional model for the lower mantle that matches that of the upper mantle for major elements is most compatible with observations and constraints, uncertainties are such that competing compositional models are tenable. Based on chondritic models, more than 90% of the mass for the Earth is composed of Fe, O, Mg and Si and the addition of Ni, Ca, Al and S accounts for more than 98% by mass the composition of the Earth. Constraints on the absolute and relative abundances of volatile elements in the Earth are consistent with only ~2% by mass of sulfur and a negligible role for H, C or N in the core.
Article
The Earth's core is about ten per cent less dense than pure iron (Fe), suggesting that it contains light elements as well as iron. Modelling of core formation at high pressure (around 40-60 gigapascals) and high temperature (about 3,500 kelvin) in a deep magma ocean predicts that both silicon (Si) and oxygen (O) are among the impurities in the liquid outer core. However, only the binary systems Fe-Si and Fe-O have been studied in detail at high pressures, and little is known about the compositional evolution of the Fe-Si-O ternary alloy under core conditions. Here we performed melting experiments on liquid Fe-Si-O alloy at core pressures in a laser-heated diamond-anvil cell. Our results demonstrate that the liquidus field of silicon dioxide (SiO2) is unexpectedly wide at the iron-rich portion of the Fe-Si-O ternary, such that an initial Fe-Si-O core crystallizes SiO2 as it cools. If crystallization proceeds on top of the core, the buoyancy released should have been more than sufficient to power core convection and a dynamo, in spite of high thermal conductivity, from as early on as the Hadean eon. SiO2 saturation also sets limits on silicon and oxygen concentrations in the present-day outer core.
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Earth's core is less dense than iron, and therefore it must contain "light elements," such as S, Si, O, or C. We use ab initio molecular dynamics to calculate the density and bulk sound velocity in liquid metal alloys at the pressure and temperature conditions of Earth's outer core. We compare the velocity and density for any composition in the (Fe-Ni, C, O, Si, S) system to radial seismological models and find a range of compositional models that fit the seismological data. We find no oxygen-free composition that fits the seismological data, and therefore our results indicate that oxygen is always required in the outer core. An oxygen-rich core is a strong indication of high-pressure and high-temperature conditions of core differentiation in a deep magma ocean with an FeO concentration (oxygen fugacity) higher than that of the present-day mantle.
Chapter
This chapter is a review of experimental results on melting temperatures of iron and Fe-rich alloys at core conditions that can thus be used to infer core temperatures. Due to its extreme importance to the understanding of the thermal structure of the core, the melting behavior of iron at high pressures has been investigated by many research groups using all of the techniques discussed in the chapter. Melting in Fe-rich iron-silicon alloys has been the subject of several previous studies using X-ray diffuse scattering, laser power-temperature discontinuities, and morphology of recovered samples. Future studies aiming to further the understanding of the temperature of the core should focus on how melting temperatures of Fe-rich systems vary with pressure and composition, especially in ternary systems, and on the Gruneisen parameters of (liquid) Fe-rich alloys at core conditions.
Chapter
This chapter develops comprehensive models of geotherm from the lower mantle (LM) to the outer core (OC) and thermal conductivity of the LM using the high-pressure and high-temperature physical properties of major constituent minerals calculated based on the ab inito computation techniques. Based on the models, the thermal property of the core-mantle boundary (CMB) region, including CMB heat flow is discussed. The CMB heat flow, which is determined by the temperature gradient and conductivity in the thermal boundary layer (TBL), controls the deep Earth dynamics, such as the lower mantle basal heating, core cooling, and dynamo activity. Analyses of the temperature profiles in the LM and OC indicate quite different CMB temperatures at the mantle and core sides. Effects of iron incorporation, which were considered indirectly in the modeling, should be clarified in the future through direct calculations under pressure with reasonable treatments.
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In the giant-impact hypothesis for lunar origin, the Moon accreted from an equatorial circum-terrestrial disk; however, the current lunar orbital inclination of five degrees requires a subsequent dynamical process that is still unclear. In addition, the giant-impact theory has been challenged by the Moon's unexpectedly Earth-like isotopic composition. Here we show that tidal dissipation due to lunar obliquity was an important effect during the Moon's tidal evolution, and the lunar inclination in the past must have been very large, defying theoretical explanations. We present a tidal evolution model starting with the Moon in an equatorial orbit around an initially fast-spinning, high-obliquity Earth, which is a probable outcome of giant impacts. Using numerical modelling, we show that the solar perturbations on the Moon's orbit naturally induce a large lunar inclination and remove angular momentum from the Earth-Moon system. Our tidal evolution model supports recent high-angular-momentum, giant-impact scenarios to explain the Moon's isotopic composition and provides a new pathway to reach Earth's climatically favourable low obliquity.
Article
An open system evolutionary model of the Earth, comprising continental crust (CC), upper and lower mantle (UM, LM), and an additional isolated reservoir (IR) has been developed to study the isotopic evolution of the silicate Earth. The model is solved numerically at 1 Myr time steps over 4.55 Gyr of Earth history to reproduce both the present–day concentrations and isotope ratios of key radioactive decay systems (Rb–Sr, Sm–Nd, and U–Th–Pb) in these terrestrial reservoirs. Various crustal growth scenarios –continuous versus episodic and early versus late crustal growth– and their effect on the evolution of Sr–Nd–Pb isotope systematics in the silicate reservoirs have been evaluated. Modeling results where the present–day UM is ∼60% of the total mantle mass and a lower mantle that is non–primitive reproduce the estimated geochemical composition and isotope ratios in Earth’s silicate reservoirs. The isotopic evolution of the silicate Earth is strongly affected by the mode of crustal growth; only an exponential crustal growth pattern with crustal growth since the early Archean satisfactorily explains the chemical and isotopic evolution of the crust–mantle system and accounts for the so-called Pb paradoxes. Assuming that the OIB source is located in the deeper mantle, our model could, however, not reproduce its target ɛNd of +4.6 for the LM, which has been estimated from the average isotope ratios of 32 individual ocean island localities. Hence, either mantle plumes sample the LM in a non–representative way, or the simplified model set-up does not capture the full complexity of Earth’s lower mantle (Nd isotope) evolution. Compared to the results obtained for a 4.55 Ga Earth, a model assuming a protracted U-Pb evolution of silicate Earth by ca. 100 Myr reproduces a slightly better fit for the Pb isotope ratios in Earth’s silicate reservoirs. One notable feature of successful models is the early depletion of incompatible elements (as well as rapid decrease in Th/U) in the UM within the initial 500 Myr, as a result of early formation of CC, which supports other evidence in favor of the presence of Hadean continental crust. Therefore, a chondritic Th/U ratio (4 ± 0.2) in the UM until 2 Gyr appears rather unlikely. We find that the κ conundrum –the observation that measured Th/U ratios and those deduced from ²⁰⁸Pb-²⁰⁶Pb isotope systematics differ– is a natural outcome of an open system evolution in which preferential recycling of U for the past 2 Gyr has played a dominant role. Overall, our simulations strongly favor exponential crustal growth, starting in the early Hadean, the transient preservation of compositionally distinct mantle reservoirs over billion year time periods, and a generally less incompatible element depleted, but non–primitive composition of the lower mantle.
Article
Recent palaeomagnetic observations report the existence of a magnetic field on Earth that is at least 3.45 billion years old. Compositional buoyancy caused by inner-core growth is the primary driver of Earth's present-day geodynamo, but the inner core is too young to explain the existence of a magnetic field before about one billion years ago. Theoretical models propose that the exsolution of magnesium oxide-the major constituent of Earth's mantle-from the core provided a major source of the energy required to drive an early dynamo, but experimental evidence for the incorporation of mantle components into the core has been lacking. Indeed, terrestrial core formation occurred in the early molten Earth by gravitational segregation of immiscible metal and silicate melts, transporting iron-loving (siderophile) elements from the silicate mantle to the metallic core and leaving rock-loving (lithophile) mantle components behind. Here we present experiments showing that magnesium oxide dissolves in core-forming iron melt at very high temperatures. Using core-formation models, we show that extreme events during Earth's accretion (such as the Moon-forming giant impact) could have contributed large amounts of magnesium to the early core. As the core subsequently cooled, exsolution of buoyant magnesium oxide would have taken place at the core-mantle boundary, generating a substantial amount of gravitational energy as a result of compositional buoyancy. This amount of energy is comparable to, if not more than, that produced by inner-core growth, resolving the conundrum posed by the existence of an ancient magnetic field prior to the formation of the inner core.
Article
The siderophile, or iron-loving elements have many applications in the Earth and planetary sciences. In primitive meteorites, differences in the relative abundances of these elements are likely due to both nebular and parent body processes. In addition, some siderophile elements are also characterised by isotopically distinctive nucleosynthetic signatures. Thus, the relative abundances and isotopic compositions of these elements can be used to trace the genetics of primary planetary building blocks. Although these elements are largely concentrated in the metallic cores of differentiated planetary bodies, their absolute and relative abundances, as well as their isotopic compositions can also reveal important information regarding conditions of core formation and the chemical evolution of the silicate portions of the planetary bodies. The lithophile-siderophile nature of the radiogenic Hf- W system allow it to be used to place chronologic constraints on planetary core formation. The differing incompatibilities of the two elements in silicate systems further mean that the system can also be used to study early differentiation processes and subsequent efficiency of mixing in the silicate portions of differentiated bodies, including Earth. The abundances of siderophile elements in the terrestrial mantle are used to assess primary and secondary melting processes, and resulting metasomatic interactions. In addition, the Re-Os isotope system can, in some instances, be used to place chronologic constraints on when these processes occurred. The abundances of siderophile elements, and Os/ Os and Os/ Os ratios in the mantle sources of ocean island basalts can be used to place constraints on the age of recycled materials, and in some instances, the types of recycled materials present in these mantle domains.
Article