Eighth International Conference on Mars (2014)
1266.pdf
CARBON RESERVOIR HISTORY OF MARS IMPLIED BY THE STABLE ISOTOPIC SIGNATURE IN
THE MARTIAN ATMOSPHERE. R. Hu1,2,3, D. M. Kass1, B. L. Ehlmann1,2, and Y. L. Yung2, 1Jet Propulsion
Laboratory, California Institute of Technology, Pasadena, CA 91109, 2Division of Geological and Planetary Sciences, California Institute of Technology, Pasadena, CA 91125, 3Hubble Fellow, Email: [email protected].
the dominant processes of atmospheric escape are nonthermal processes.
The two most important non-thermal escape processes are pick-up ion sputtering and photochemical
processes. The rates of both processes have been studied extensively [6] and an emergent trend in the past
decade is the application of 3-D Monte Carlo direct
simulations [14,15], which we adopt as the nominal
values. The photochemical loss rates have been calculated for present-day solar EUV conditions [15], and
the appropriate way to scale up the rates with the solar
EUV flux to 3.8 Ga is unclear. We thus explore the
effects of scaling the photochemical loss rate with the
solar EUV flux from linear to cubical.
The fractionation factor of CO photodissociation, a
dominant photochemical process of producing escaping carbon [15], has never been evaluated for Mars.
Here we estimate this factor, by considering that the
energy from the incident photon, in excess of the bond
dissociation energy, is imparted to carbon and oxygen
atoms as kinetic energy (Figure 1). We find that the
fractionation ratio 12C/13C via CO photodissociation is
0.6 for both the current and early Sun. This fractionation factor is very significant, due to two effects: first,
the conservation of momentum determines that 13C
takes a lesser fraction of the excess energy than 12C in
each photodissociation event; second, 13C has a greater
critical energy to escape from the gravity of Mars.
1
12
C
C
0.9
13
0.8
Normalized Production Rate
Introduction: The evolution of the atmosphere of
Mars is one of the most intriguing problems in the exploration of the Solar System. Presently Mars has a
very thin atmosphere in equilibrium with polar caps
and regolith. Yet, the climate of Mars during the Noachian and Hesperian Era could have been significantly
warmer than the present [1]. From the early state to the
present state, Mars may have experienced major climate change.
Since CO2 is the major constituent of Mars’s atmosphere, its isotopic signatures offer a unique window to the evolution of climate on Mars. The Sample
Analysis at Mars (SAM) instrument on the Mars Science Laboratory (MSL) has reported the most precise
measurement of the carbon isotopic signature of CO2
in the martian atmosphere: δ13C = 46±4 per mil [2].
Motivated by this new measurement, we develop a
carbon reservoir and evolution model to trace the history of δ13C of Mars’s atmosphere. The carbon history
of Mars has been extensively studied with reservoir
models [3-7]. Here we adopt the model of [4] but make
major changes in accordance with recent progress of
our understanding of Mars. Specifically, we use the
latest estimates of the non-thermal escape rates, and
constrain our models with observational evidence of
carbonate deposition on Mars.
Carbonate Reservoir of Mars: Carbonate has
been found to be a minor component of the martian
soil by the Phoenix Lander [8] and MSL [9] but not at
other landing sites. Based on these measurements, we
estimate that the upper 1 m of soil can store up to ~ 1
mbar of CO2 as carbonate (i.e., ~2 wt. %). This amount
is considered to be the upper limit of carbonate formation during the Amazonian Era. We consider the
carbonate formation rates during the Noachian and
Hesperian Era to be a free parameter, as formation of
carbonates may have been widespread on Mars. Carbonate-bearing rocks have been discovered in various
Noachian terrains [10-12]. MSL could potentially assess Hesperian carbonates at Gale crater and provide
further constraints to our models.
Processes of Atmospheric Escape: We model
carbon reservoir and 13C history starting from 3.8 Ga,
i.e., the end of the early/mid Noachian and the Late
Heavy Bombardment (LHB). Before then, the atmosphere on Mars would be close to a steady-state balance
between volcanic outgassing, hydrodynamic loss, and
impact removal and delivery of volatiles [13]. After,
0.7
0.6
0.5
0.4
0.3
0.2
0.1
0
0
0.5
1
1.5
Kinetic Energy of Carbon [eV]
2
2.5
Figure 1: Energy distribution of carbon atoms produced by
photodissociation of CO, in comparison with the critical
energy for each isotope to escape (dashed lines). This is calculated using the current solar minimum spectrum for an
exobase at 200 km and recently measured branching ratios
[16]. Using early Sun proxies [17] or assuming higher exobases give quanlitatively similar results.
Eighth International Conference on Mars (2014)
High early carbonate formation
Pressure [Bar]
10
Low early carbonate formation
Equivalent pressure of free carbon including the atmosphere,
the absorbed carbon in the regolith, and the polar caps
1
0.1
0.01
40
δ 13C
20
0
−20
−40
−60
10
Rate [Bar/Gyr]
Results: With the newly calculated fractionation
factor for CO photodissociation, we find that the photochemical processes are highly efficient in enriching
δ13C and thus carbonate formation is required to compensate this effect. Prelimenary studies have found the
following two classes of potential solutions assuming
the nominal values of the sputtering rates [14].
High early carbonate formation: If the photochemical loss rates scale as the cube of the solar EUV flux,
δ13C produced by atmospheric escape would be greater
than 100 per mil without carbonate formation. Therefore, very efficient carbonate formation must have occurred before 3.0 Ga to yield the correct present-day
δ13C (Figure 2). How much early carbonate formation
is required depends on the amount of late carbonate
formation. For a maximum of 1 mbar carbonate
formed during the Amazonian, we find that the minimum amount of early carbonate formation would be
equivalent to 5 bar of CO2. This scenario would imply
a thick atmosphere on early Mars. However, this scenario, requiring in excess of 10 wt. % carbonate in the
ancient crust, is inconsistent with present knowledge of
the geologic record [10-12].
Low Early Carbonate Formation: We consider another scenario with reduced escape rates. If the photochemical loss rates only scale as the square of the solar
EUV flux, the minimum amount of carbonate formation before 3.0 Ga would be equivalent to 0.4 bar of
CO2 (Figure 2). Although sputtering is the dominant
non-thermal escape process, photochemical loss significantly contributes to the 13C enrichment. The amount
of carbonate, ~1 wt. % globally averaged in the ancient
crust, is consistent with the geologic record.
Discussion: Both classes of potential solutions
suggest that substantial amounts of carbonate have
formed during the late Noachian and Hesperian. The
scenario with low early carbonate formation suggests
the atmosphere had a surface pressure on the order of a
few 100 mbar before 3.0 Ga, sufficient to cause transient melting on the surface [18]. Alternatively, stronger
greenhouse gases, such as CH4 and H2, might have
contributed to warming the early Mars’s atmosphere
[19,20].
Our result highlight the crucial importance of a reliable understanding of non-thermal escape processes
on Mars. Whether early Mars had a thick atmosphere
or a thin one, is highly sensitive to the details of the
atmospheric escape during the Amazonian. How the
photochemical loss rates scale with the solar EUV flux
plays a key role, and is yet to be studied theoretically.
The Mars Atmosphere and Volatile EvolutioN
(MAVEN) mission will provide data to calibrate current non-thermal escape models and improve the extrapolation to early Mars.
1266.pdf
Carbonate formation rate
1
0.1
0.01
0.001
0.0001
1
1.5
2
2.5
3
Time [Gyr]
3.5
4
4.5
5
Figure 2: Scenarios of carbon reservoir evolution on Mars
since the LHB.
References: [1] Fasset C. I. and Head J. W. (2011) Icarus, 211,
1204. [2] Webster C. R. et al. (2013) Science, 341, 260. [3] Jakosky
B. M., et al. (1994) Icarus, 111, 271. [4] Kass D. M. (1999) PhD
Thesis. [5] Jakosky B. M. and Phillips R. J. (2001) Nature, 412, 237.
[6] Chassefiere E. and Leblanc F. (2004) P&SS, 52, 1039. [7] Gillmann C. et al. (2009) EPSL, 277, 384. [8] Sutter B. (2012) Icarus,
218, 290. [9] Leshin L. A. et al. (2013) Science, 341, 6153. [10]
Ehlmann B. L. et al. (2008) Science, 322, 1828. [11] Morris R. V. et
al. (2010) Science, 329, 421. [12] Niles P. B. et al. (2013) SSR, 174,
301. [13] Lammer H. et al. (2013) SSR, 174, 113. [14] Leblanc F.
and Johnson R. E. (2002) JGR, 107, 5010. [15] Groeller H. et al.
(2014) P&SS, in press. [16] Gao H. et al. (2013) JPCA, 117, 6185.
[17] Ribas I. and Miralda-Escude J. (2007) A&A, 464, 779. [18]
Forget F. et al. (2013) Icarus, 222, 81. [19] Wetzel D. T. et al.
(2013) PNAS, 110, 8010. [20] Ramirez R. M. et al. (2014) Nature
Geo, 7, 59.