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レビューノート\\
Hashimoto, G.L., Y. Abe, and S. Sugita (2007)
The chemical composition of the early terrestrial atmosphere: Formation of a reducing atmosphere from CI-like material
J. Geophys. Res., 112, E05010, doi:10.1029/2006JE002844.


\section*{要旨}

The chemical composition of Earth’s early atmosphere is likely to have
played an important role in the origin of life. In particular, the redox
state of Earth’s early atmosphere may have controlled greatly the
efficiency of abiotic synthesis of biologically important organic
compounds. However, the chemical composition of the Earth’s early
atmosphere has not been studied extensively before. In this study, we
theoretically estimate the chemical composition of an atmosphere near
the end of the completion of planetesimal accretion of the Earth. Our
calculation results show that the Earth’s early atmosphere is very
reducing, rich in H2 and/or CH4, regardless of which meteoritic
materials accrete and regardless of which chemical reactions control the
molecular abundance in the atmosphere. In other words, impact-degassed
origin of Earth’s atmosphere leads inevitably to a reducing chemical
composition, which is very favorable to efficient prebiotic synthesis of
organic matter on Earth.

\section{Introduction}

惑星形成の標準モデルは, 微惑星の集積 
(Safronov, 1972; Hayashi et al., 1985).
微惑星集積過程では, 重力エネルギーの解放により原始惑星を加熱
(原始惑星が成長する程加熱は大きくなる).
原始惑星がある大きさに達すると衝突脱ガスにより惑星をまとう
(
Lange and Ahrens, 1982; 
Ahrens et al. ,1989).


地球大気の成長過程のごく初期では非常に還元的(H$_2$ rich)であると
考えられる. 理由は 2 つ.
\begin{itemize}
 \item 微惑星の衝突により金属鉄が常に供給され, 
金属鉄が大気の化学組成をコントールするから.
    \begin{itemize}
     \item Dreibus and Wanke, 1989
     \item Abe et al., ¨ 2000
    \end{itemize}
 \item おそらく nebula gas の取り込みもあるだろう.
    \begin{itemize}
     \item Porcelli and Pepin, 2000
    \end{itemize}
\end{itemize}
ごく初期では, 厚い水蒸気大気が形成されると考えられる.
\begin{itemize}
 \item Abe and Matsui, 1986
 \item Matsui and Abe, 1986
 \item Zahnle et al., 1988
\end{itemize}


しかし, 集積がほとんど完了すると, 集積によるエネルギー供給が減少, 
原始海洋が形成
[e.g., Abe and Matsui, 1986; Matsui and Abe, 1986]. 
原始海洋は large impact により何度も蒸発しただろう
[e.g., Sleep et al., 1989]. 
38 億年前(重爆撃が終了する頃)以降に, おそらく表層は生命が発生できる
環境になったのだろう
[e.g., Maher and Stevenson, 1988; Kasting, 1993 and references
therein]. 
しかし, この時期の大気組成はわかってない.
この論文ではそれを考える.

[5] When most of H2O in the steam atmosphere collapses to form the
ocean, a large amount of H2 will be left in the atmosphere. It is not
evident whether H2 is lost through hydrodynamic escape in a short time
or not [e.g., Sekiya et al., 1980, 1981; Tian et al., 2005]. However,
the atmospheric composition was probably not controlled by chemical
reaction with metallic iron because the supply of metallic iron is
greatly reduced and metallic core segregation is also likely to have
finished by this time (i.e., probably 4.5 billion years ago) [e.g., Abe,
1993; Abe et al., 2000]. Thus the dominant contributor to the atmosphere
during the final stage of Earth’s accretion is most likely volatile
degassing upon meteoritic impacts. 


地球集積の最終段階では, 酸化をする物質が集積したと
推定される.
なぜなら, マントルには親鉄元素(Pt, Pd, Rh, Ru, Os, Ir, Au, Re)
が多いから
[e.g., Wanke et al., 1984]. 
マントル中の親鉄元素量は, 金属鉄との平衡状態における値よりも
ずっと大きい.
This excess of highly siderophile elements in the mantle
can be accounted for by accretion of oxidizing material, such as CI
chondrite, after the end of core formation [e.g., Kimura et al., 1974;
Chou et al., 1983]. Such a latestage accretion of oxidizing material,
so-called late veneer, is generally believed to generate an oxidizing
atmosphere that is composed largely of H2O and CO2. 
An atmosphere
dominated by CO2-N2-H2O, which is often called a neutral atmosphere, is
referred to as an oxidizing atmosphere in this paper in order to clarify
the contrast to a reducing atmosphere. However, whether H2O and CO2 were
the main components of gas phase degassed from CI chondrites is not
examined yet.

[7] The composition of Earth’s early atmosphere has been also
considered rather oxidizing: CO2-rich mixtures with only trace amount of
H2 and CH4 [e.g., Rubey, 1955; Kasting, 1993]. This prevailing view is
based on the fact that modern volcanic gas composition is relatively
oxidizing [e.g., Holland, 1984]. Such oxidizing volcanic gases are
consistent with the composition predicted for equilibrium with their
mantle sources that have O2 fugacities of the quartz-fayalite-magnetite
(QFM) buffer [Holland, 1984]. After the removal of metallic iron from
the Earth’s mantle as a result of early core formation [e.g.,
Stevenson, 1990], the upper mantle has been near the QFM-buffered redox
state [e.g., Delano, 2001]. Thus volcanic gases released from Earth’s
interior after iron segregation should have been relatively oxidizing
[Holland, 1984]. 

[8] However, a reducing early atmosphere is favored by researchers
studying the origin of life. Here we use the term ‘‘reducing
atmosphere’’ to refer to an atmosphere containing H2, CH4, and CO as
major components. Experiments on the prebiotic synthesis indicate that
the oxidation state of the atmosphere controls the yields of the
synthesis of organic matter greatly [e.g., Miller, 1953; Pinto et al.,
1980; Stribling and Miller, 1987]. A reducing atmosphere, which contains
reducing species, such as CH4, H2, and CO, allows efficient abiotic
synthesis of organic compounds, but it is very inefficient to synthesize
the organic matter in the atmosphere which does not contain reducing
species. 

[9]
In the following, we first discuss the chemical composition
of impact-degassed atmosphere. Second, we consider an influence of
preexisting atmosphere that is formed during the main stage of accretion
on the chemical composition of the early atmosphere. Third, we infer the
chemical composition of the early atmosphere from the elemental
abundance of the present Earth surface, and we discuss the life span of
the early reducing atmosphere.

\section{Late-Veneer Atmosphere}

Accretion of oxidizing material after the end of core formation
(late veneer) has been suggested as a mechanism to supply not only
excess siderophile elements in the Earth’s mantle but also atmophile
elements in the Earth’s near surface [e.g., Chyba, 1990]. 
\footnote{
Nature 343, 129-133 (11 January 1990) | doi:10.1038/343129a0; Accepted 20 November 1989

Impact delivery and erosion of planetary oceans in the early inner Solar System

Christopher F. Chyba
}
It is very
likely that atmophile elements in the accreting material of late veneer
will degass upon impact since partial vaporization of impactor will
occur when the proto-Earth is larger than 75\% of the present mass [e.g.,
Abe, 1993]. 

[11] At the time of later-veneer accretion, the average temperature of
Earth’s surface would not be high enough to allow chemical reactions to
proceed efficiently since accretionary flux at that time is too small to
sustain a massive
steam atmosphere [e.g., Abe, 1993]. However, materials involved in an
impact vapor plume will be in chemical equilibrium immediately after the
impact since the rates of chemical reactions are fast under
high-temperature and highpressure conditions of the impact vapor
plume. Because an impact vapor plume cools as it expands, chemical
equilibrium becomes increasingly more difficult to achieve. The chemical
reactions of vapor plumes are quenched when reaction rates become too
slow to attain chemical equilibrium. 

\subsection{Accreting Material}

[12] To
examine the atmospheric composition generated by the late veneer, we
assume that the bulk composition of planetesimal accreted as the late
veneer is similar to that of CI-chondritic material. It is widely
accepted that the composition of the late-veneer material is similar to
CI chondrite [e.g., Chou et al., 1983]. 

[13] The composition of CI
chondrite has been studied extensively, and its chemical composition is
mostly established [e.g., Wasson and Kallemeyn, 1988]. However, the
abundance of water is still controversial [e.g., Kuramoto and Matsui,
1996 and references therein]. Some groups of researchers found about 20
wt\% of water in CI chondrite [Wiik, 1956], whereas 5 10 wt% of water has
also been reported [Boato, 1954; Kerridge, 1985]. The variation in water
abundance in CI chondrite may be related to the contamination of
terrestrial water [Kaplan, 1971] although it is still
controversial. Thus we consider two cases: dry case $6 wt% H2O, and wet
case $20 wt\% H2O (Table 1) in this study. 

[14] To evaluate the
atmospheric composition generated by the late veneer, we have to take
into account not only atmophile elements but also iron and
sulfur. Although most of the major atmophile elements would remain in
the atmosphere, the atmospheric abundance of oxygen is controlled by the
partition between fluid (i.e., gas and liquid) phases and solid
phase. The change in the redox state of iron affects the abundance of
oxygen in the atmosphere. Also, the abundance of iron, which can combine
with oxygen, is determined by the abundance of sulfur combined with iron
since iron combined with sulfur does not combine with oxygen. 

[15] In
this study, we calculate thermochemical equilibrium of a system that
consists of six elements (H, C, N, O, S, and Fe). Since Mn, Co, Ni, and
Cu are similar to Fe in chemical behavior, the abundances of these
elements are added to that of Fe. Other major elements are removed from
the reaction system as oxides because their affinity to oxygen is much
stronger than Fe. To remove these elements together with their combining
oxygen, the following valences are assumed: Na(+1), Mg(+2), Al(+3),
Si(+4), P(+5), K(+1), Ca(+2), Ti(+4), and Cr(+3) [Railsback,
2003]. Minor elements not mentioned here are neglected in our
calculation. 

[16] The elemental abundances used in our calculation are
summarized in Table 1. The chemical composition of CI chondrite given by
Wasson and Kallemeyn [1988] is used to estimate the elemental abundance
of the wet case. The H2O content of the wet case corresponds to $\sim$ 18
wt\%. We assume that the H2O content of dry case is 5.8 wt\% as reported
by Kerridge [1985]. The elemental abundance of the dry case is created
by removing some H2O from the composition of the wet case.

\begin{table}
%
%Component O Element H C N O S Fe
%a
%
%Component R 4.000 1
%
%Wet Case 1.219 0.164 0.007 1 0.113 0.215
%
%Dry Case 0.612 0.291 0.012 1 0.201 0.380
%
\caption{Table 1. Model Composition of Two Components.
All values listed are in mole and are normalized to O  1.
} 
\end{table}

\subsection{Calculation of Chemical Equilibrium}

[17] To examine the atmospheric composition generated by the late
veneer, we have calculated the equilibrium chemical composition at a
variety of temperatures and pressures, ranging from 300 to 2500 K and 1
to 100 bar, respectively. The conditions of quench of chemical reactions
in the impact vapor plumes are determined not only by reaction rates but
also by the cooling rate of vapor plume. Although several studies
estimated the effective quenching temperature to be higher than 1000 K
for impact vapor plumes [Fegley et al., 1986; Zahnle, 1990], it depends
on various parameters, such as size of impactor and impact
velocity. Although a kinetic model that involves all the reaction paths
relevant to each elemental composition is necessary for precise
determination of the molecular composition, equilibrium molecular
composition at different P-T conditions gives us a general view of the
reaction system and serves as the baseline of more complicated studies. 

[18] To obtain the molecular composition at equilibrium, we employed a
 thermodynamic energy minimization procedure that considers 30 species
 (Table 2) and determines the complete equilibrium composition of a
 system that minimizes the total Gibbs free energy as a function of
 temperature, pressure, and elemental composition. We used the
 thermodynamic data on the JANAF Tables [Chase, 1998].

[19] It is important to note that the oxygen fugacity of the system is
not assumed but is calculated in the model. In principle, it cannot be
determined which buffer reaction controls the oxygen fugacity of the
system until the actual equilibrium calculation is conducted. Thus
procedures that assume a buffer reaction of oxygen fugacity a priori are
inappropriate for examining the chemical composition of the
atmosphere. In this study, we calculate the partitioning of oxygen
between fluid and solid phases, and the oxygen fugacity of the system is
self-consistently determined. 

\subsection{Atmospheric Composition}

[20] Figures 1 and 2 show dry gas compositions as functions of quenching
temperature. Although the fraction of water is not shown in these
figures, the equilibrium compositions of atmospheres contain a large
amount of water. The fraction of water in the gas phase was not shown
because it changes greatly with temperature via a simple phase change
(i.e., condensation and vaporization), not via chemical reactions as
other chemical species do. The amount of water released from the
accreting material depends on the quenching temperature and water
content of the accreting material. When the accreting material contains
$\sim$ 6 wt\% H2O (dry case), approximately 50 -- 100\% 
of hydrogen is released
as H2O. When the accreting
material contains $\sim$ 20 wt\% H2O (wet case), approximately 75 100\% of
hydrogen is released as H2O. 

[21] Here it is noted that the CI chondrites are generally considered to
be one of the most oxidizing materials in the solar system based on the
state of iron [e.g., Wasson, 1985]. The CI chondrites are the only clan
of chondrites, with no members of which contain metallic iron; all the
iron in the CI chondrites is oxidized [e.g., Weisberg et al.,
2006]. Thus the compositions of atmospheres obtained in the above
calculation using the CI composition would be considered as the most
oxidizing atmosphere generated by the accretion process. The most
important point here is that considerable amounts of reducing species
such as H2, CO, and CH4 are present in ‘‘CI-derived’’ atmosphere
despite of the oxidizing nature of CI composition. The abundance of
reducing species is insensitive to the pressure and the water content of
CI chondrite. Although reducing gas species are minor for very low
quenching temperatures ($\sim$ 300 K), it is likely that the quenching
temperatures for impact vapor plumes are higher than 1000 K [Fegley et
al., 1986; Zahnle, 1990]. Therefore an atmosphere generated by the
accretion of CI chondrite is likely to contain more than 30\% of reducing
species, such as H2 and CO. The atmospheric H2/CO ratio depends on the
water content of CI chondrite; the H2/CO ratio is high for high water
content. 

[22] The proportion of reducing species (for example, H2 and CO)
decreases as quenching temperature decreases from 1800 to 1400 K. The
decrease in the proportion of reducing species is caused by the
formation of iron sulfides. Although iron sulfides are unstable at high
temperatures, they are stable and formed from iron oxides in this
temperature range. The transformation of iron oxides into iron sulfides
releases oxygen from solid phases into the atmosphere. Since this
released oxygen combines with reducing species, the proportion of
reducing gas species in the atmosphere decreases.



Table 2. Chemical Compounds and Their Phases Considered in This Study
Name Gas Phase Hydrogen Water Oxygen Carbon monoxide Carbon dioxide Methane Nitrogen Ammonia Sulfur Sulfur Mercapto Hydrogen sulfide Sulfur oxide Sulfur dioxide Sulfur oxide Carbon oxide sulfide Iron sulfide Iron Hydrogen Hydroxyl Condensed Phase Water (liquid) Graphite Pyrite Troilite Iron sulfide (liquid) Iron (crystal) Iron (liquid) Wustite ¨ Magnetite Hematite H2O C FeS2 FeS FeS Fe Fe Fe0.947O Fe3O4 Fe2O3 H2 H2O O2 CO CO2 CH4 N2 NH3 S S2 SH H2 S SO SO2 S2O COS FeS Fe H OH Chemical Formula



Figure 1. Equilibrium composition of dry gas phase at (a) 1 bar and (b)
100 bar total pressures. Water content in CI chondrite is about 20 wt\%
(wet case). This figure does not include water vapor since atmospheric
water varies greatly due to condensation. 


[23] When quenching temperature is decreased from 1000 to 500 K, CH4 is
formed instead of H2 and CO. The mixing ratio of CH4 is more than
several percent when quenching temperature is between 400 and 900
K. Although these temperatures are lower than the expected gas-phase
quenching temperature of impact vapor plume, CH4 may be produced by
surface catalysis reaction [Sekine et al., 2003; Kress and McKay,
2004]. 

[24] If the chemical equilibrium is attained at very low
temperatures (for example, 300 K), reducing gas species that exist at
high temperature are converted into graphite. Therefore there is an
extremely large amount of reducing species in the condensed phase even
though the atmosphere is composed of oxidizing species, such as CO2 and
N2. However, no mechanism has been suggested to accelerate the
production of graphite at such low temperatures. 

[25] The range of volume mixing ratio of ammonia (NH3) calculated here
is several to several hundred parts per million. The sensitivities of
NH3 abundance to temperature and pressure are similar to those of
CH4. The abundance of NH3 is maximized in the temperature range from 600
to 900 K at the pressures used in this study (1 100 bar). The
abundance of NH3 increases with pressure. When pressure is higher than
10 bar, NH3 mixing ratio is higher than 1 ppm for a wide range of
temperature. 


Figure 2. Equilibrium composition of dry gas phase at (a) 1 bar and (b)
100 bar total pressures. Water content in CI chondrite is about 6 wt\%
(dry case). This figure does not include water vapor since atmospheric
water varies greatly due to condensation.


\subsection{Oxygen Deficit}

[26] The results of chemical equilibrium calculations demonstrate that
the Earth’s early atmosphere would contain a significant amount of
reducing species even if atmosphere is entirely generated by the
accretion of CI chondrites. The reason why this occurs is the
following. As mentioned above, all the iron in the CI chondrites is
oxidized [e.g., Wasson, 1985; Weisberg et al., 2006]. However, CI
chondrite also contains reducing components such as organic matter; the
components in CI are not in chemical equilibrium to each other, no
matter which temperature is assumed. The gases formed through the
decomposition of organic matter are intrinsically reducing. The redox
state of degassed atmosphere from CI is determined by the relative
abundance of oxygen in the oxidized iron to that of reducing organic
matter. 

[27] The deficit in oxygen is easily demonstrated by a simple
calculation used by Kasting et al. [1993]. It is appropriate to assume
that carbon content of 3.5 wt\% in CI chondrite represents the abundance
of reduced carbon in CI chondrite since almost all carbon in CI
chondrite is in reduced forms [e.g., Kerridge, 1985; Wasson and
Kallemeyn, 1988]. If reduced carbon is oxidized by conversion of ferric
iron to ferrous iron, two moles of ferric iron is required to oxidize 1
mol of reduced carbon.

\begin{eqnarray}
 Fe2 O3  C \rightarrow 2FeO  CO 1
\end{eqnarray}

Therefore about 32.6 wt\% of ferric iron is required to oxidize 3.5 wt\%
of reduced carbon. However, a typical iron content in CI chondrite is
about 20 wt\% [Wasson and Kallemeyn, 1988]. The abundance of ferric iron
in CI chondrite is not sufficient to oxidize all the carbon in CI
chondrite. Even though some of the reducing gases are oxidized through
the reaction with iron oxides, the abundance of oxygen stored in iron
oxides is not enough to oxidize all the reducing gases released from
organic matter. 

[28] It is inevitable that an atmosphere generated by
the late veneer contains a considerable amount of reducing gases since
oxygen is deficient even in CI chondrites, which are one of the most
oxidizing materials among the primitive materials in the solar
system. Thus the formation of reducing atmospheres by accretion is a
very robust consequence. Even when chemical equilibrium is not attained
in the vapor plume, the atmosphere will contain reducing species as a
result of oxygen deficit.


\section{The Influence of the Preexisting Atmosphere}

[29] It is not obvious whether the very early atmosphere and ocean
formed before the accretion of the late veneer are completely lost or
not [Hashimoto and Abe, 1995]. Here the late veneer is defined as the
mass that is not reacted with metallic iron, so the preexisting
atmosphere (including ocean) prior to the later-veneer accretion is the
atmosphere that was reacted with metallic iron. In this section, we
examine the influence of survived preexisting atmosphere
on the chemical composition of the early atmosphere at the end of the
late-veneer accretion. 


\subsection{Atmospheres Before the Late Veneer}

[30] The formation of atmosphere by accretion of volatilecontaining
planetesimals must have begun in the early stage of Earth formation
[e.g., Abe et al., 2000]. Atmophile elements degas from accreting
planetesimals, when impact generates sufficiently high-temperature and
high-pressure conditions. Shock experiments indicate that degassing
begins when the mass of the proto-Earth is as large as 10\% of the
present size [Lange and Ahrens, 1982; Ahrens et al., 1989]. It is
consistent with argon isotope data which suggest that most of the
terrestrial atmosphere was formed by 4.0 billion years ago [e.g., Hamano
and Ozima, 1978]. 

[31] The atmosphere formed during the main
accretionary phase is probably reducing since a large amount of metallic
iron, which later becomes the core of the Earth, would be contained in
the planetesimals [e.g., Dreibus and Wanke, ¨ 1989]. Then, a large
amount of H2 would be formed during planetary accretion by the reaction
of metallic iron with ¨ H2O [e.g., Dreibus and Wanke, 1989]. Thus the
atmosphere formed during the main accretionary phase is expected to be
reducing even when the atmosphere is generated solely by degassing. 

[32]
If Earth accreted in the solar nebula, hydrogen-rich atmosphere also
could have been formed by gravitational capture of solar nebula gas
[e.g., Hayashi et al., 1979]. Although solar nebula gas cannot be a
major source of the present terrestrial atmosphere, it has been
suggested that some part of captured solar nebula gas might be retained
as a source of rare gases [e.g., Porcelli and Pepin, 2000]. Since H2 in
a solar-type atmosphere will react with Fe-bearing silicate on the
planetary surface [e.g., Sasaki, 1990], the composition of the resulting
atmosphere is similar to that of the atmosphere generated by degassing
during the main accretionary phase. 

[33] There are several processes
that may dissipate a hydrogen-rich atmosphere formed during the main
accretionary phase. Hydrogen in the upper atmosphere is heated by strong
radiation of X ray and extreme ultraviolet (XUV) from the young Sun and
is lost hydrodynamically to the space; i.e., hydrodynamic escape [e.g.,
Sekiya et al., 1980, 1981]. This process was believed to dissipate
efficiently in a short time (about several million years) a
hydrogen-rich atmosphere. However, it has been suggested that an escape
of hydrogen from early Earth’s atmosphere would occur much slower than
previously thought [Tian et al., 2005]. 

[34] Another mechanism that
dissipates the atmosphere is impact erosion [e.g., Melosh and Vickery,
1989; Ahrens, 1993]. Impact erosion of atmosphere would reduce the
relative contribution of the preexisting atmosphere to the current
terrestrial atmosphere. However, the degree of the impact erosion of
Earth’s atmosphere throughout its history is difficult to estimate
accurately. There are a couple of important factors controlling the
efficiency of the atmospheric erosion by impacts. First, the large mass
of Earth prevents this process from occurring efficiently. It requires
impact velocities higher than twice the escape velocity [Vickery and
Melosh, 1990; Melosh et al., 1993], but such an extremely high velocity
impact is expected to be rare [e.g., Chyba, 1991]. Second, the condition
of Earth’s surface at the
time of large impacts also changes the erosion efficiency greatly. When
Earth is covered with an ocean, the erosion efficiency is much greater
than the case when Earth is covered with a dense steam atmosphere [Genda
and Abe, 2003, 2005]. Thus, although an extremely large impact, such as
moon-forming giant impact, may remove a significant fraction of the
preexisting atmosphere, the fraction of the lost atmosphere to the total
mass of the atmosphere due to such a giant impact depends on the
condition of Earth’s surface. Despite these uncertainties, Earth’s
atmosphere during and after the late-veneer accretion phase is predicted
to be reducing as we discuss in the following sections 3.2 and
3.3. 

\subsection{Two-Component Model for Atmospheric Composition }


[35] In order
to consider the influence of the preexisting atmosphere formed before
the accretion of late veneer, we use a model in which the composition of
the early atmosphere is described as mixtures of two chemically
different components, R and O. Since metallic iron has a predominant
influence on the elemental abundance of atmosphere, we choose the
component R to be the component equilibrated with molten metallic iron
and the component O to be the volatile directly degassed from accreting
planetesimals during the late stage not equilibrated with metallic
iron. 

[36] The mixing ratio of the two components f is defined as the
fraction of the late-veneer contribution to the volatile inventory in
the Earth surface and mantle (i.e., atmosphere, ocean, crust, and
mantle). Although f can be defined by using any volatile element,
hydrogen is chosen in this study.  

\begin{eqnarray}
  f \equiv O NH O NH 
\end{eqnarray}

where NO is the number of hydrogen supplied by the H component O, and NR
is that supplied by the component R. H If the atmosphere and ocean
generated during the main accretionary phase are completely dissipated
after the end of accretion of core-forming metallic iron, the atmosphere
is generated entirely by the late veneer. Then, the atmosphere does not
contain any component that is equilibrated with molten metallic iron, f
value is unity. In contrast, f value is less than unity if the very
early atmosphere formed during the accretion of core-forming metallic
iron has not been completely lost. 

[37] The elemental abundance in the component R is determined by
equilibrium between the atmosphere and molten metallic iron. When the
atmosphere and metallic iron are in equilibrium, oxygen fugacity is
buffered by iron oxidation reaction. The abundances of carbon and
nitrogen in the gas phase are also controlled by the solubility of these
elements in molten metallic iron. 

[38] We assume that the component R
consists of equal amounts of H2 and H2O (Table 1). When metallic iron
and silicates exist simultaneously, oxygen fugacity would be close to
that of the iron-wu ¨stite buffer [Holland, 1984; Abe et al., 2000]:

\begin{eqnarray}
  2xFe  O2  = 2Fex O
\end{eqnarray}

where x = 0.947. The ratio of H2/H2O under the condition of iron-wustite
buffer will be larger than unity at high ¨ temperatures (i.e., >1000 K)
[Sasaki, 1990]. In a real system, where other species coexist, the
activity of wustite ¨ is nearly equal to the molar fraction of FeO in
silicate melts, and oxygen fugacity is possibly smaller than that
controlled by the iron-wu ¨stite buffer. Thus the assumed ratio of
H2/H2O = 1 is an underestimate for a component-R atmosphere; a real
component-R atmosphere may be more reducing than our model. 

[39] We also
assume that carbon and nitrogen are not present in the component-R
atmosphere since solubilities of carbon and nitrogen in molten metallic
iron are much larger than that of hydrogen [e.g., Wood, 1993]. However,
some carbon and nitrogen would remain in the atmosphere. The influence
of these remaining carbon and nitrogen in the component R is discussed
below (section 3.4). 

\subsection{Molecular Composition of Mixed Atmosphere}

[40]
The preexisting atmosphere will chemically interact with the material
accreted later since vapor plumes generated by impacts of planetesimals
interact intensely with the surrounding atmosphere during their
downrange motion and buoyancy uprise [Sugita and Schultz,
2003]. Although such high-temperature conditions are local and
transient, such impacts occur frequently at the final stage of the Earth
accretion. Since it is difficult to estimate the extent of interaction
between the preexisting atmosphere and accreting material, we examined
an extreme case that shows that the materials accreted later are
chemically equilibrated with the preexisting atmosphere. This situation
will occur when repeated impacts efficiently promote the chemical
reaction between the accreting material and the preexisting
atmosphere. 

[41] Using the CI chondrite composition for the component O,
we calculated the chemical composition of mixed atmosphere (Figures 3
and 4). Since the value of f represents the fraction of the
metallic-iron-free material accreted, the abundance of reducing species
decreases as the value of f increases. When quenching temperature is
higher than $\sim$ 1000 K, the abundance of H2 increases with f due to the
survival of H2-H2O atmosphere. However, the abundance of CH4 increases
significantly when quenching temperature is lower than $\sim$ 800 K. 
At low
temperatures, hydrogen in the H2-H2O atmosphere combines with carbon
delivered by late-stage accretion. If some catalytic reactions can
produce CH4 at low temperatures [e.g., Sekine et al., 2003, 2006], the
survival of H2-H2O atmosphere probably leads to CH4rich atmosphere at
the end of accretion. 

\subsection{Estimation of $f$ Value }

[42] The fraction f of
the late-veneer contribution (i.e., component O) can be estimated from
the relative abundance of atmophile elements. There are marked
differences in siderophility among atmophile elements. The atmosphere of
the component R is depleted in siderophile elements since it has been
removed by dissolving in metallic iron. However, the atmosphere that
originated from the component O is not depleted in siderophile
elements. Therefore we can evaluate f, the fraction of late-veneer
contribution, using two atmophile elements with different
siderophility. 

[43] We choose hydrogen as a representative of weakly siderophile elements, and we select carbon and nitrogen for
those of strongly siderophile elements. If metallic iron absorbed all of
the carbon and the nitrogen in the atmosphere of the component R but no
hydrogen (see section 3.2), then the f value is given by 

\begin{eqnarray}
 f = O Earth NH NX O Earth NX NH 4
\end{eqnarray}

where NO and NEarth are the numbers of ith element (i = H, i i X; X = C,
N) in the component O and the Earth surface and mantle,
respectively. The derivation of equation (4) is given in Appendix A. 



Figure 3. Dry gas composition for various values of parameter f. Water
content of CI chondrite is about 20 wt\% (wet case). Total pressure and
temperature of each system are (a) P = 1 bar, T = 500 K, (b) P = 1 bar,
T = 800 K, (c) P = 1 bar, T = 1000 K, and (d) P = 1 bar, T = 1500
K. This figure does not include water vapor since atmospheric water
varies greatly due to condensation.


[44] To evaluate f value, we use the elemental abundance in the Earth
surface and mantle given in Table 3 [Ronov and Yaroshevsky, 1976;
Holland, 1978; Kargel and Lewis, 1993]. Here we consider the Earth
surface as the sum of the atmosphere, the ocean, and the crust. We also
take into account the hydrogen and carbon in the mantle since water and
carbon would be exchanged between the crust and the mantle [e.g.,
Kasting and Holm, 1992; Sleep and Zahnle, 2001]. Using the composition
of component O given in Table 1, we calculated the values of f (Table
4). The calculated f values for wet CI composition and dry CI
composition are about 0.3 and 0.1, respectively, which are significantly
lower than unity. In other words, despite the model uncertainties, the
present elemental abundance of the Earth surface and mantle requires a
significant contribution from the atmosphere that was generated when
metallic iron was present.

[45] As discussed above in section 3.2, the value of f calculated by
equation (4) is an overestimate since some amounts of carbon and
nitrogen remain in the atmosphere of component R and some hydrogen
components dissolve in molten metallic iron. When we take this
overestimation of f into account, a larger amount of the atmosphere
generated under the presence of metallic iron (i.e., before and during
core formation) needs to remain in
the present terrestrial atmosphere than our calculated f values
indicate. 


Figure 4. Dry gas composition for various values of parameter f. Water
content in CI chondrite is about 6 wt\% (dry case). Total pressure and
temperature of each system are (a) P = 1 bar, T = 300 K, (b) P = 1 bar,
T = 800 K, (c) P = 1 bar, T = 1300 K, and (d) P = 1 bar, T = 1500
K. This figure does not include water vapor since atmospheric water
varies greatly due to condensation. 



[46] It is also likely that the f value of the early
atmosphere is much smaller than that estimated from the present
elemental abundance because hydrogen in the early atmosphere was likely
lost by escape. The elemental abundance might be also altered by
exchange of atmophile elements between the mantle and the core. Although
it is difficult to evaluate precisely the influence of the exchange
between the mantle and the core, its influence is likely to be much
smaller than the effect of hydrogen escape. Thus the hydrogen in the
surface layers of the Earth must have decreased over the course of its
evolution; the ratios of hydrogen to carbon/ nitrogen would be much
higher in the past (see section 4.1). 


\section{Discussion and Implication}

\subsection{The Life Span of Reducing Atmosphere}

[47] A reducing early atmosphere is indicated by the elemental abundance
of atmophile elements in the Earth
surface and mantle. Such a
reducing atmosphere needs to have evolved to an oxidizing atmosphere
found today. It is likely that the early reducing atmosphere is oxidized
through hydrogen escape [e.g., Catling et al., 2001; Tian et al.,
2005]. Since hydrogen is the lightest element, it is preferentially lost
to the space. Preferential loss of hydrogen relative to other atmophile
elements increases the relative abundance of oxygen in the
atmosphere. When hydrogen photodissociated from water vapor is lost, the
remaining oxygen, a by-product of photodissociation, would oxidize the
atmosphere. 

[48] A small value of f indicates that most of the hydrogen
in the Earth’s surface are delivered during the main accretionary
phase. The atmosphere formed during planetary accretion (component-R
atmosphere) would contain a large amount of H2 as a result of the
reaction of metallic iron with H2O. Since such an early atmosphere would
contain more H2 than H2O in mole [e.g., Dreibus and Wanke, 1989; Sasaki,
1990; Abe et al., 2000], the ¨
amount of dissipated hydrogen must be more than that in the present
 terrestrial ocean. In addition to the dissipation of H2, there is a
 need to oxidize the reduced carbon, such as C, CO, and CH4. There is
 also probably a need for oxidizing the mantle since ferric iron in the
 mantle would be produced by oxidation of ferrous iron. These oxidation
 processes will consume a great number of oxygen atom [Kasting et al.,
 1993] and require more hydrogen to escape.



Table 3. Distribution of Carbon, Nitrogen, and Hydrogen in the Earth
Surface and Mantlea Hydrogen Atmosphere Hydrosphere Crust Mantle Earth
Surface Earth Surface + Mantle a

Carbon 5.8 3.3 1.0 1.2     10 1018 1022 1022
16

Nitrogen 2.8 1.6 1.4 1.1  1020  1018  1020  1020

1.6  1023 6.7  1022 1.5  1023 2.2  1023 3.7  1023

1.0  1022 2.2  1022

4.3  1020 5.4  1020

All values listed are in mole.



 [49] The timescale for hydrogen escape, which is the timescale for
 oxidization of the reducing atmosphere, is controlled by the energy
 available to drive the escape flow since escape of hydrogen from a
 hydrogen-rich atmosphere is energy-limited [Tian et al., 2005]. Since
 hydrogen in the upper atmosphere is heated by radiation of XUV from the
 young Sun, the upper limit of mass loss rate is controlled by the solar
 XUV flux. 

[50] The escape of hydrogen from very early Earth’s
 atmosphere might occur at a very high rate since the XUV emission of
 the young Sun is estimated to be about 100 times that of the present
 Sun [Ribas et al., 2005]. However, the hydrogen escape depends greatly
 on whether the Earth was surrounded by a protoplanetary disk
 gas. Before the protoplanetary disk gas was dissipated, the XUV
 radiation was absorbed by the disk gas and could not have a great
 influence on the planetary atmosphere. When once the disk gas was
 dissipated, the strong radiation would start to affect the planetary
 atmosphere. The timing of disk gas dissipation is not observationally
 constrained yet. However, theoretical models of terminal phase of
 terrestrial planet formation require the presence of thin disk gas
 around the terrestrial planets; high eccentricities of terrestrial
 planets resulting from the final stage of planet formation through
 mutual collision of Mars-sized protoplanets (i.e., giant impacts) are
 very difficult to reduce without remnant protoplanetary disk gas
 [Kominami and Ida, 2004; Nagasawa et al., 2005]. 

[51] If a massive hydrogen escape occurs on very early Earth, it might change the composition of preexisting atmosphere significantly. If accretion of metallic iron was ended and metallic iron disappeared from the Earth surface before the dissipation of the protoplanetary disk gas, the strong solar XUV radiation would deplete the atmospheric H2 from the preexisting atmosphere. Although some H2O would be lost by photodissociation of H2O and resulting hydrogen would escape from the Earth, most of the H2O likely remains on the Earth because of the cold trap, which effectively prevents a massive loss of H2O.

[52] If a massive hydrogen escape generates a H2O-rich preexisting
atmosphere, the atmospheric composition subsequent to the late-veneer
accretion would be similar to that of a case for f = 1 regardless of the
actual value of f. As shown in section 2, an atmosphere with f = 1 has
the most oxidizing composition among the atmosphere predicted for the
early Earth, but it is still very reducing and rich in H2 and/or CH4. It
is worth noting that such a H2O-rich preexisting atmosphere does not
lead to oxidation of the atmosphere generated by the late-veneer
accretion since H2O does not have excess oxygen to oxidize the reducing
species. Thus, even in an extreme case where H2 in the preexisting
atmosphere is completely lost and only H2O is left behind on the Earth
before the accretion of the late veneer, the resulting atmosphere after
the late veneer accretion would be reducing (i.e., rich in CH4 and/or
H2). 

[53] It is noted that a reducing atmosphere at the time of
late-veneer accretion would likely survive more than a few hundred
million years. When the heavy bombardment of planetesimals is almost
ended about 3.8 billion years ago, the solar XUV flux decreased to about
10 times the present value [Ribas et al., 2005]. On the basis of the
calculation made by Tian et al. [2005], the hydrogen escape flux would
be about 6  1015 molecules m2 s1 for a solar XUV level 10 times that of
today. Then, dissipation of hydrogen equivalent to that in the
terrestrial ocean needs 500 million years. Thus it is likely that
ancient terrestrial atmosphere had contained reducing species at least
for a few hundred million years. 

\subsection{Implication for Early Earth}

[54]
The formation of a reducing early atmosphere does not depend on the
reaction kinetics since the reducing early atmosphere is formed as a
result of the deficiency of oxygen. The result of our model indicates
that the early atmosphere would contain a significant amount of reducing
species, such as H2, CO, CH4, and NH3. Such a reducing atmosphere would
lead to efficient abiotic synthesis of organic compounds for the origin
of life. 

[55] It has also been suggested that reducing gases, such as
CH4 and NH3, have been important greenhouse gases on the early Earth
[e.g., Sagan and Mullen, 1972; Kiehl and Dickinson, 1987; Sagan and
Chyba, 1997; Pavlov et al., 2000]. The geologic record tells us that
liquid water was present on the early Earth [e.g., Mojzsis et al., 2001;
Wilde et al., 2001]. However, standard models of solar evolution
indicate that the luminosity of the Sun was $\sim$ 70\% 
of the current value at
the time when the Sun is formed and has
increased monotonically [e.g., Gough, 1981]. To compensate for the
fainter solar luminosity, strong greenhouse effect is required on the
early Earth to raise the surface temperature to above the freezing point
of water [e.g., Kasting and Toon, 1989]. 


Table 4. Estimated Value of f From the Elemental Abundance Near the Earth Surface
Water Content H/C Wet case Dry case H/N Wet case Dry case Reservoir Surface Surface Surface Surface Surface Surface Surface Surface + mantle + mantle + mantle + mantle f Value 0.34 0.44 0.10 0.13 0.34 0.25 0.10 0.07



[56] It would be an important
future work to estimate the actual mixing ratio of relevant molecules,
such as CH4 and NH3. It is necessary to use a modeling of chemical
kinetics to estimate the precise chemical composition consequent on the
planetesimal impact since temperature and pressure dependences of each
reaction path are different. Chemical interaction between the
preexisting atmosphere and the accreting material also deserves further
study since the extent of interaction between impact vapor plume and
surrounding atmosphere does affect the atmospheric composition,
especially the mixing ratio of CH4. 

\subsection{Implication for Venus and Mars}

[57] It is likely that the early atmospheres of Venus and Mars are also
reducing similarly to that of the Earth since the atmospheres on both
planets would also be generated by degassing of planetesimals. 

[58] The
formation process of the Venusian atmosphere would be similar to that of
the Earth’s atmosphere since the mass and the heliocentric distance of
Venus are similar to the Earth. Then, the Venusian atmosphere should
also evolve from a reducing one into the present oxidizing one. Hydrogen
escape must be an essential process for evolution of the Venusian
atmosphere since the present Venusian atmosphere contains almost no
hydrogen [e.g., Kasting and Pollack, 1983; Kasting, 1988]. 

[59] If the
Venusian CO2 atmosphere is generated from CI chondrite, the amount of
hydrogen accreted is as large as 2.3 8.1  1019 kg. That is 0.15  0.52
times of the mass of hydrogen in the terrestrial ocean. A hydrogen
escape rate estimated for a solar XUV level 10 times the present based
on the calculation made by Tian et al. [2005] indicates that it takes 40
130 million years to dissipate hydrogen from Venus. When f value is not
unity, the timescale for hydrogen escape is 1/f times longer (see
Appendix B). If the f value for the Venusian atmosphere is similar to
that of the terrestrial atmosphere, it takes more than a billion years
to dissipate hydrogen from the Venusian atmosphere. 

[60] It would be
worth mentioning that nonthermal escape processes might significantly
enhance the rate of hydrogen escape on Venus since Venus could have lost
its magnetic field very early in its history. It has been also suggested
that the solar wind may have been much more massive in the distant past
[Wood et al., 2002]. Without the protection of magnetic field, very
strong solar wind would probably have caused strong hydrogen escape in
addition to the hydrodynamic escape although the details of nonthermal
escape processes need to be worked on. 

[61] The early atmosphere of Mars would also contain reducing gases
although formation process of the Martian atmosphere may be somewhat
different from those on Earth and Venus. The mass of Mars, a tenth of
the mass of the Earth, is rather small for impact vaporization to occur,
but it is large enough to cause impact degassing and impact melting
[e.g., Abe, 1993]. Chemical reactions would proceed in a
high-temperature region near impact points [Kuramoto and Matsui, 1996].

[62] Reducing gases in the Martian early atmosphere could have been
important in warming the surface of early Mars since fluvial features on
the Martian surface indicate that surface condition of early Mars was
warm and wet [e.g., Carr, 1996]. Radiative warming by reducing gases,
such as CH4 and NH3, will be favorable for providing the necessary
greenhouse warming since it is difficult to warm early Mars by the
greenhouse effect of a dense CO2 atmosphere [Kasting, 1991].

\section{Conclusion}

[63] The Earth’s early atmosphere is likely to contain a significant
 amount of reducing species, such as H2, CO, and CH4, when the accretion
 is mostly completed. This conclusion is robust because reducing
 atmosphere is generated even if atmosphere is not contributed by
 captured solar nebula gas or H2 generated by reaction with metallic
 iron and is solely generated by accretion of CI chondrite. Since CI
 chondrites are the most oxidizing kind of meteorites, the atmosphere
 generated by CI chondrites will be the most oxidizing primitive
 atmosphere that is allowed in the currently accepted planetary
 formation theories. 

[64] The elemental abundance of the present Earth
 surface and mantle strongly suggests that most of atmophile elements in
 the present terrestrial atmosphere were once equilibrated with metallic
 iron. Since the ratios of hydrogen to carbon and nitrogen in the Earth
 surface and mantle are higher than those of the probable carrier (i.e.,
 CI-like material) of Earth’s volatile components, mechanism(s) that
 change the relative abundance of atmophile elements are
 needed. Dissolution of carbon and nitrogen into the molten metallic
 iron during the main accretionary phase is the most probable process
 that can account for the elemental abundance of the Earth.

[65] The most plausible process that changes the early reducing
atmosphere into the present oxidizing one is hydrogen escape. When
hydrogen photodissociated from water vapor is lost, the remaining
oxygen, a by-product of photodissociation, would oxidize the
atmosphere. It is indicated that early atmosphere is rich in hydrogen
since this oxidation process requires a large amount of hydrogen to
escape. If escape flux of hydrogen is energylimited, the timescale of
hydrogen escape would be as long as about a billion years. Thus, once a
reducing atmosphere is created, it may exist for about a billion years
on the early Earth.

\appendix 

\section{Derivation of Equation (4)}

[66] Since NH and NH are the numbers of hydrogen supplied by the
components R and O, respectively, the number of hydrogen in the Earth
surface and mantle, N Earth, H is given by

\begin{eqnarray}
 N^{Earth}_{H} = R O NH  NH  NH
\end{eqnarray}

while the number of carbon or nitrogen in the Earth surface and mantle
is given by

\begin{eqnarray}
  Earth O NX  NX
\end{eqnarray}

since we assumed that carbon and nitrogen are not present in a
component-R atmosphere. Using these two equations, we can reformulate
equation (2) as

\begin{eqnarray}
 f \equiv O NH O NH
R NH
O NH Earth NH
O Earth NH NX Earth N O NH X  O  Earth  NH NX  O Earth NX NH
\end{eqnarray}

\section{Timescale for Hydrogen Escape}

[67] When escape flux is energy-limited, the timescale for hydrogen
escape is in proportion to the amount of hydrogen accreted. We can
reformulate equation (A3) and obtain an expression for the amount of
hydrogen accreted in terms of the amount of carbon accreted and the
value of f: Earth NH

\begin{eqnarray}
O NH O NC
  1 Earth NC f
\end{eqnarray}
 This indicates that the amount of hydrogen accreted is inversely
 proportional to the value of f, when the amount of carbon accreted is
 some fixed value.


\end{document}
