# Interior Controls on the Habitability of Rocky Planets

Cedric Gillmann  Kaustubh Hakim  KU Leuven, Institute of
Astronomy,Celestijnenlaan 200D, Leuven, 3001, Belgium  Royal Observatory of
Belgium,Ringlaan 3, Brussels, 1180, Belgium  Diogo Lourenço  ETH Zürich,
Institute for Geophysics,Sonneggstrasse 5, Zürich, 8093, Switzerland  Sascha
P. Quanz  ETH Zürich, Institute for Particle Physics and
Astrophysics,aWolfgang-Pauli-Str. 27, Zürich, 8093, Switzerland  Paolo A.
Sossi  ETH Zürich, Institute for Geochemistry and Petrology, Clausiusstrasse
25, Zürich, 8092, Switzerland

###### Abstract

No matter how fascinating and exotic other terrestrial planets are revealed to
be, nothing generates more excitement than announcements regarding their
habitability. From the observation of Mars to present-day efforts toward Venus
and the characterization of exoplanets, the search for life, or at least
environments that could accommodate life, has been a major drive for space
exploration. So far, we have found no other unquestionably habitable world
besides Earth. The conditions of the habitability of terrestrial planets have
proved elusive, as surface conditions depend on the complex interplay of many
processes throughout the evolution of a planet. Here, we review how the
interior of a rocky planet can drive the evolution of surface conditions and
the atmosphere. Instead of listing criteria assumed to be critical for life,
we discuss how the bulk-silicate planet can affect the onset, continuation and
cessation of habitability. We then consider how it can be observed and current
efforts towards this end.

##  1 Introduction

Investigations into habitability have provided ample opportunity for
reflection as to its nature, but also planetary evolution processes in
general. Agreeing on a definition for habitability has provedtroublesome.
Rather than providing a definition, we consider that the concept can be
distilled to a simple question: can life develop there? Unfortunately, the
notion of life is heavily biased toward life as we know it on Earth, and even
the question of ”development” remains vague: is survival enough? Does it
include its proliferation instead? Does it cover the emergence of life? Would
the requirements not change with time, and because of life itself? The
interested reader could consult the review by [1] for a discussion of the
definition of the term.

It is nonetheless possible to draw a list of requirements that any living
organism we can imagine would need to remain active. Those would include (i) a
solvent (water, on Earth), (ii) energy and (iii) the presence of various
elements and nutrients in sufficient abundances and forms to facilitate (i)
and (ii). It has been argued that energy was likely easily accessible in
general, from solar radiation, redox potential or other processes, and that
life as we know it makes use of some of the most common building blocks
(elements) in the universe, the CHONPS. Even minor elements that are used by
life, such as Fe or Mg, for instance, are relatively abundant on the surface
of the Earth (even if those may vary from one specific organism to the other).
Finally, the solvent common to known life, water, is one of the most abundant
molecules in the Solar System.

Indeed, spectroscopic surveys show that compositions of stars in the solar
neighbourhood are similar to that of our Sun [2]. However, despite the
ubiquity of the right elemental building blocks, the universe does not appear
to be teeming with life. This indicates the need for the appropriate physico-
chemical conditions to arise on planetary surfaces. For example, water would
need to be liquid, which imposes further constraints on the extant pressure
and temperature conditions.

There are also important differences between conditions required for the
emergence of life and its continuation: no credible model for abiogenesis has
been proposed, we are still unable to produce new life through laboratory
experiments and have not recorded spontaneous life formation [3, 4].

Many specific characteristics have been proposed as essential to planetary
habitability [[, see]]cockell2016. Some are related a planet’s galactic
location, its star and orbit [5], the pattern of the other planets in the
system or the presence of a large moon [6]. Other are related to the planet’s
properties like its composition [7], inherited from its origins (see section
2), its size/mass (see Section 3) or processes [[, such as plate tectonics,
for example]]foley2015role.

However, it has been suggested that piling up these criteria will inherently
result in a rare-Earth-type situation [9] that only an infinitesimal fraction
of terrestrial planets could fulfil [10]. In short, precise lists describe
only Earth. While it is possible that Earth could be the template for all
habitable worlds, and that those are incredibly rare, limiting our approach
leaves gaps in our understanding of life in general.

In the end, based on the above, if a single criterion is to be kept, it should
be the presence of liquid water. Disregarding individual organisms
characteristics, in terms of biology, liquid water is deemed necessary both
for the emergence and persistence of life [[, e.g.,]]dartnell2011, and
therefore roughly demarcates a range of surface conditions amenable to the
occurrence of life. Investigating how a planet can reach and maintain
conditions compatible with the presence of liquid water remains the best
general-purpose criterion we can apply.

This idea was the seed for the development of the concept of habitable zone
[[, see]for example, but references to this concept date back to at least the
19th century]lammer2009,lingam2021. The circumstellar (abiotic) habitable zone
[[, HZ]]callejas2014 is defined as the shell of space around a star where
liquid water can be maintained at the surface of a planet. Its inner boundary
is set by the maximum energy flux that can be received in the presence of
liquid water (see Section 4). Its outer boundary is set where the maximum
greenhouse can no longer sustain temperatures high enough for liquid water
[[]]kopparapu2013. Most depictions of the HZ assume an Earth-like planet with
an Earth-like atmosphere, highlighting the importance of atmospheric
composition.

A mix of N2, CO2 and H2O [[]]kasting1993 is often considered, but other gases
would contribute, like CH4 or H2 [[, e.g.]]pierrehumbert2011. The HZ also
varies with the type of star considered and its age, as its luminosity changes
throughout its evolution [[, e.g.,]]rugheimer2013,rugheimer2018.

This simple definition has historically been used to select targets that are
deemed likely to be habitable (see, e.g. [20]).

A more complete definition would consider the whole planet as a dynamic
system, possibly including life itself. It remains uncertain how such extended
definitions of habitable zones could be translated into practical parameters
or observables, but future observation and characterization of exoplanets will
help refine this discussion.

The presence of liquid water implies sufficiently clement conditions that H2O
(l) is more stable than its gaseous equivalent, water vapour. Therefore, the
habitability of a terrestrial planet, depends on the nature of its atmosphere.
Hence, processes that affect the volatile inventory (such as the accretion
history or atmospheric escape) or the atmosphere response to solar radiation
(radiative transfer, the greenhouse effect or atmosphere dynamics) influence
planetary habitability. However, as one of the drivers of long-term planetary
evolution, the state of the interior of terrestrial planets is critical to the
origin and evolution of the atmosphere and surface conditions. The dynamics of
the mantles of planets provides a link between the atmosphere/surface and the
deep interior, which extends down to the core [21, 22]. Mantle thermal
history, dynamics and eventually melting cause volcanism and volatile delivery
to the atmosphere. Interior dynamics also play a part in the regulation of the
climate on Earth, by recycling surface material and volatiles into the
interior. This cycle is deeply intertwined with outgassing, weathering and
subduction. On Earth, it notably affects the CO2 concentration inn the
atmosphere and is named long-term carbon cycle. Likewise, the thermal
evolution of the core governs the generation of a magnetic field that affects
atmospheric escape processes and long-term atmospheric evolution.

In short, a comprehensive picture of planetary evolution requires
consideration of the planet as a complex system, that depends on the
efficiency of numerous feedback mechanisms. Here, we focus on aspects relative
to the interior of rocky planets and only include succinct descriptions of the
interactions between those mechanisms and the interior of planets for clarity.
The present review focuses on the relationship between habitability and the
silicate planet. Initially, and in order to set the boundary conditions from
which a given rocky planet is thought to have evolved, the molten state of the
mantle is considered, before transitioning to a solid state, and its long-term
interactions with the surface and atmosphere of the planet.

Habitability has been described in the context of targeted studies of specific
planets like Mars [23] or Venus [24], or in terms of a descriptive approach
detailing each mechanism that can affect habitability and the atmosphere [[,
e.g.]]gillmann2022. Here, we instead draw from the core science questions that
remain unsolved, summarise the current state of our knowledge and discuss how
future work can build upon this foundation. We have identified four broad
questions that are the pillars of ongoing investigations into the habitability
of terrestrial planets in the Solar System and beyond:

  * •

How do planets become habitable and to what degree is it influenced by their
accretion history, bulk composition and initial conditions?

  * •

How do the interior dynamics of a planet and its mantle convection regime
affect habitability?

  * •

What can cause habitability to end or push a planet out of the regulating
cycles that maintain habitable surface conditions?

##  2 The onset of habitability

###  2.1 Volatile Budgets

Both planetesimal- and pebble accretion imply a considerable degree of
stochasticity as to the volatile element budgets of growing terrestrial
planets [25, 26]. Planets accumulate material from a wide range of
heliocentric feeding zones [27], regardless of whether condensed material was
initially confined to an annulus [28] or ring [29, 30] or whether it resulted
from inward-drifting pebbles [31]. This stochasticity leads to a gradual
decline in elemental abundance with increasing volatility (decreasing
condensation temperature111The temperature at which half of the mass of a
given element condenses from a gas of solar nebular composition, at a nominal
pressure of 10-4 bar.) in the terrestrial planets [26]. Elements that were
sequestered into the core during its formation (siderophile elements) show
depletions in the mantle with respect to the ‘volatility trend’ defined by
silicate-loving (lithophile) elements [[, e.g.,]]fegley2020volatile.
Therefore, although the total budgets of siderophile elements in the
terrestrial planets are higher than those observed in their mantles, they are
inaccessible to geochemical scrutiny, and do not play a direct role in
determining habitability conditions at the surface of the planet [[, although
they may indirectly contribute, such as in the generation of a
geodynamo,]]nimmo2004influence.

As such, the bulk planetary abundances of key atmosphere-forming moderately
volatile elements, of which S and P are part, can be estimated from those of
lithophile elements of similar volatility elements. Conversely, highly
volatile elements, notably C, H and N, because of their low abundances in
planetary mantles with respect to the nebular gas or chondritic meteorites [[,
e.g.,]]mccubbinbarnes2019, are likely to have been delivered irregularly and
near the tail end of accretion [[, e.g.,]]albarede2009,marty2012. As a result,
their provenance is likely distinct from that of major, planet-forming
elements [[, see]]dauphas2017, broadley_etal2022, such that predicting their
abundances in the terrestrial planets is fraught with uncertainty.

Through analysis of peridotitic rocks that comprise the Earth’s mantle,
together with inversion of the compositions of products of mantle melting in
the ocean basins , the concentrations of volatile elements in the convecting
mantle of the Earth can be determined [[,
cf.]]marty2012,palmeoneill2014,hirschmann2018. The mass of oxygen is
overwhelmingly contained in the Earth’s mantle, bound in silicates and oxides,
and can be computed by stoichiometry to better than 2 % relative. The
moderately volatile elements, P and S, also reside chiefly in the mantle, and
have abundances that can be determined to a precision of $\sim$20 % relative
[40, 38].

, H, C and N are presumed to exist as a particular species (H2O, CO2, N2)
during mantle melting, and their abundances in basalts are normalised to those
of a refractory lithopile element (H2O/Ce CO2/Ba or CO2/Nb), or failing that,
another volatile element (Ar for N2) that behaves in a similar manner [[,
e.g.,]]hirschmann2018, with uncertainties of 50 % relative. Furthermore,
significant fractions of the budgets of H, C and N are present on the surface
[[, atmosphere + oceans + crust;]]rubey1951 of the Earth, known to within 20 %
relative. [[, see]and references therein]hirschmann2018. Together, the
abundances of H, C and N in the bulk silicate Earth (BSE) can be computed
within a factor $\sim$2\. With these caveats in mind, the depletion factors of
the six key elements, CHONPS, in the BSE relative to the solar- and CI-
chondritic compositions are found in Table 1.

Table 1: The abundances of the life-essential volatile elements, C, H, O, N, P and S, in the bulk silicate Earth (BSE), CI chondrites and in the Sun. All concentrations expressed in weight fraction. Uncertainties for C, H and N in the BSE are of the order of a factor $\sim$2, while P and S are $\pm$20 %, with O $\pm$2 % | C | H | O | N | P | S
---|---|---|---|---|---|---
BSE |  1.4$\times$10-4 |  1.2$\times$10-4 | 0.4433 |  2.8$\times$10-6 |  8.7$\times$10-5 |  2.0$\times$10-4
CI chondrite | 0.0348 | 0.0197 | 0.459 |  2.95$\times$10-3 |  9.85$\times$10-4 | 0.0535
Sun |  2.17$\times$10-3 | 0.738 |  6.34$\times$10-3 |  7.48$\times$10-4 |  6.44$\times$10-6 |  3.42$\times$10-4
Depletion (/CI, Al) | 0.0014 | 0.0022 | 0.341 |  3.3$\times$10-4 | 0.031 | 0.0013
Depletion (/Sun, Al) |  1.56$\times$10-4 |  3.93$\times$10-7 | 0.169 |  9.04$\times$10-6 | 0.033 | 0.0014

A striking observation is that no more than $\sim$0.4 ppm (by weight, ppmw) of
solar nebula matter remains in the Earth, based on the Al-normalised abundance
of H (Table 1). Of course, this estimate ignores any amount of H potentially
dissolved in the core. Assuming a partition coefficient between core and
mantle, D${}^{H}_{core/mantle}$ = 10 [42], then the 120 ppmw H in the mantle
would give 0.12 wt. % in the core, consistent with geophysical constraints
[43]. Even if D${}^{H}_{core/mantle}$ = 100 [[, as permitted by the
computations of]]li_etal2020Hcore and 1.2 wt. % H were to be dissolved in the
core, the depletion factor of H relative to Al and the solar nebula (Table 1)
would be $\sim$10-5, lower than or equivalent to, respectively, the solar-,
Al-normalised abundances of C or N, even without accounting for the
siderophile behaviour of the latter pair [[,
e.g.,]]speelmanns_etal2019,grewal2022,blanchard_etal2022metal.

Therefore, it seems that, compared to the Sun, the Earth (or its building
blocks) underwent preferential depletion of H relative to similarly volatile
elements, C and N, with the noble gases having been depleted by an even
greater factor than H [47, 48]. Indeed, [47] recognised that “the earth has
only one millionth part of its proper quota of inert gases”, with modern
estimates [35, 49] ranging from 1 part in 10 billion (Ne) to one one-millionth
(Xe). Hence, the Earth has retained but a minuscule fraction of a solar
component (i.e., from the nebular gas).

The origin of the small fraction of volatiles has been reconciled with either
i) dissolution and subsequent loss of a component of the nebular gas [[,
e.g.,]]olson2019nebular or ii) predominantly dry accretion of the Earth with a
small component of volatile-rich, chondritic and/or cometary material [[,
e.g.,]]albarede2009. Isotope ratios among volatile elements offer a
complementary way to distinguish between these plausible sources [[, see]for a
review]peron2018origin. The isotopic compositions of noble gases (except He
and Ne), H, C and N in the BSE overlap with those of chondrites [52, 53, 37],
which, themselves, are distinct from that of the solar nebula (as inferred
from the Sun and the atmospheres of the gas giants).

Models of dissolution of the solar nebula can account for the observed
abundance of H relative to the noble gases [54, 55]. [54] highlight the
plurality of sources that could have been accreted to the Earth to satisfy
both elemental and isotopic constraints, with N and heavy noble gas abundances
being consistent with chondritic and cometary sources, respectively, whereas H
and the light noble gases permit a greater fraction of nebular material. Yet,
such a process would have resulted in an overabundance relative to the BSE
(Table 1) and an H isotope composition far too light (D/H of $\sim$30 $\times$
10-6), necessitating isotopic fractionation associated with its subsequent
loss via atmospheric escape. These models have, so far, not considered carbon
in their treatment, making it unclear whether dissolution of the nebular gas
is capable of producing a match to the HCN abundances of the BSE.

Figure 1: The abundances of the life-essential volatile elements, CHONPS,
relative to the solar abundances [56] normalised by Al in: CI chondrites [[,
teal line,]]palmeoneill2014, CO chondrites [[, orange
line,]]alexander_etal2012 and the bulk silicate Earth [[, BSE; green
line,]]marty2012,palmeoneill2014,hirschmann2018. The dashed orange line
represents 2.5 % of the CO abundances.

On the other hand, to first order, the relative abundances of volatile
elements bear a strong resemblance to those in chondrites [[]Fig.
1]broadley_etal2022, suggesting that the present-day budgets of these elements
were brought by such materials [53, 57], although Ne isotopes permit a small
fraction of its budget in the BSE to be solar in origin [58, 59, 54]. However,
it should be kept in mind that this is relative to the 10-7 that remains of
the Ne budget of the solar nebula.

The precise nature of the chondritic material that could have delivered these
volatiles, and, what fraction of their total budgets were contributed by such
material, remains under scrutiny. Assuming that the D/H and N isotopic
compositions of the BSE reflect those of their sources, the Earth’s H and N
(and, by extension, C) budgets were thought to have been carried by CI- or CM-
type carbonaceous chondrites [[, e.g.,]]alexander_etal2012,furimarty2015. This
interpretation has been called into question on the grounds that the Earth is
known to share an isotopic heritage with enstatite chondrites (EC) for the
major rock-forming elements [[, e.g.,]]dauphas2017. Consequently, [53]
postulated that EC-like material could have delivered much of the H budget of
the Earth. A major caveat is that D/H and N isotope ratios could be
fractionated in a mass-dependent manner [45, 54], and thus not only record the
provenance of these elements, but also mass transfer processes (such as
atmospheric escape) that could have modified their terrestrial inventories
thereafter.

Figure 1 indicates that the addition of 2.5 % CO-chondritic material by mass
to an otherwise volatile-free Earth would reproduce the present-day H, C and N
budgets of the BSE, within uncertainty. Equally, 0.7 % of CI-like material,
consistent with the mass of the late veneer [[, cf.]]albarede2009, would
satisfy the H budget of the Earth, but would deliver 6 times too much N.
However, the observation that H, C and N are depleted relative to the noble
gases [[, normalised to CI chondrites;]]halliday2013 suggests that a
straightforward chondritic contribution is too simplistic, and that these
elements preserve some vestige of the main accretion phase. Therefore, in the
event that chondritic material delivered all of the H, C and N, this would
have represented, at most, $\sim$10–25 % and 2.5% of the P and O budgets of
the BSE, respectively. Hence, P and O must have delivered during the main
accretion phase of the Earth and hence from predominantly non-carbonaceous
material [[, e.g.,]]dauphas2017. The addition of 0.7 % CI or 2.5 % CO-like
material would exceed the S budget of the BSE by 1.5–2.5 $\times$. However, S,
together with Se and Te, was likely depleted by the extraction of a small
amount of sulfide-rich melt from the mantle after the cessation of core
formation, but before the accretion of the late veneer [61, 40].

In sum, the bulk volatile inventory of terrestrial planets, taking the Earth
as an example, is low ($\leq$ 0.05 % by mass). This observation, together with
the chondritic relative abundances and isotope ratios of major volatiles,
indicates chondritic material (although of uncertain origin) as the major
source of H, C and N in the BSE. Their isotopic and elemental compositions are
more difficult to reconcile with capture and dissolution of the nebular gas,
without appealing to additional processes (i.e., atmospheric escape) and would
require additional investigation. Instead, it points towards accretion via a
mixture of largely volatile-depleted bodies with smaller amounts carbonaceous
chondrites [[, though enstatite chondrites remain feasible]]piani_etal2020,
representing $\geq$30 % of the present-day BSE budgets of H, C and N [[, see
also]]steller_etal2022. Budgets of the less volatile P ($\leq$25 % CC-
material) and O ($\leq$2 % CC-material) were brought predominantly by sources
other than carbonaceous chondrites, whereas S, likely reflects the composition
of the late veneer.

###  2.2 Volatile Distribution

The constitution of an atmosphere, during a magma ocean epoch, is dictated by
the abundances of elements available in planet-building material (see section
2.1), the intensive variables, namely temperature and pressure that determine
the speciation of gaseous molecules, and the solubility relations these
gaseous species have with their dissolved counterparts in silicate liquids.
Because the masses of atmospheres around terrestrial planets are small
relative to their interiors, the assumption is frequently made that the
interior imposes the state of the atmosphere. As such, the partial pressure
(p) of given species, i, in the atmosphere is:

$p_{i}={x_{i}}\frac{gM^{a}}{4\pi r^{2}},$ (1)

where g = the acceleration due to gravity, $M^{a}$ is the mass of the
atmosphere, $x_{i}$ is the molar fraction of element i and r is the planetary
radius. To determine the value of $M^{a}_{i}$ requires knowledge of the
distribution of a given gas species in the magma ocean using a solubility law.
A solubility law relates the fugacity of species i, fi, (or partial pressure,
given an ideal gas expected to typify most secondary atmospheres around
terrestrial planets) to the mole fraction, X, of the relevant species
dissolved in a condensed phase, here, a silicate liquid;

$X_{i}=\alpha f_{i}^{\beta},$ (2)

where $\alpha$ is a Henrian constant and $\beta$ is the stoichiometric
coefficient. The relationship between the fraction of mass of species $i$ in
the atmosphere (M${}^{a}_{i}$) relative to that in the mantle (M${}^{m}_{i}$),
assuming that the sum of the two is equivalent to the total mass,
M${}_{i}^{T}$ = M${}^{a}_{i}$ \+ M${}^{m}_{i}$, and that fi = pi with $\beta$
= 1, is given by:

$\frac{M^{a}_{i}}{M^{m}_{i}}=\frac{4\pi r^{2}}{gM_{T}\alpha},$ (3)

where $M_{T}$ is the total mass of the planet. Taking a generic chemical
reaction and setting $\alpha$ = 1000 ppm/bar (in SI units, 10-8 Pa-1) and
$\beta$ = 1, the distribution of a volatile species, comprising 1000 ppm of
the planet, between the mantle and atmosphere is determined as a function of
planetary mass, $M_{T}$, where g is given at known r by solving the Adams-
Williamson equation of state [[, see]]Valencia:2007jv,sossi_etal2023.

Figure 2 shows that Eq. 3 predicts that the quantity
M${}^{a}_{i}$/M${}^{m}_{i}$ decreases as planetary mass (proportional to g)
increases. That is, as shown by [64], small, rocky bodies will have a greater
fraction of a given volatile element present in the atmosphere relative to the
molten interior as compared to larger planets. Despite the decreasing relative
fraction of a given volatile element in the atmosphere, its partial pressure
(or fugacity) in the atmosphere increases asymptotically to a maximum value
given by Eq. 2; $f_{i}$ = $X_{i}$/$\alpha$. Consequently, mass is a key
property that dictates whether the planet is capable of harbouring an
atmosphere.

Figure 2: The distribution of a fictive element, i, with $\alpha$ = 1000
ppm/bar and $\beta$ = 1 (cf. eq. 2) between the molten interior (m) and
atmosphere (a), expressed in terms of its mass Mi and fugacity (in bar) as a
function of the log10 of the planetary mass in kg. The planet is assumed to be
Earth-like, in which an silicate mantle comprises 70 % of its mass, with the
remaining 30 % being made up of an Fe-rich core.

Although some species, notably N2 [[, at high fO2]]libourel_etal2003nitrogen,
CO2 [66], H2O [[, at high fH2O;]]stolper1982water and H2 [[, at high
fH2;]]hirschmann_etal2012 dissolve in silicate melts with $\beta$ $\sim$ 1,
others have more complicated dissolution stoichiometries. At low $f$H2O
($\leq$ 1000 bar), water (and possibly H2 at low fH2), dissolves into silicate
melts with $\beta$ = 0.5 [69, 64]. At low fO2 nitrogen dissolves in silicate
melts as nitride, N3-, a species that is more soluble than the diatomic
molecule [65, 70]. The solubility of sulfur has long been known to be redox-
sensitive [71, 72, 73], existing in silicate liquids as S2- at low fO2 and
SO${}_{4}^{2-}$ at high fO2.

In this context, the oxygen fugacity (fO2) holds special significance, as,
unlike the other CHONPS elements, O is invariably in excess in the mantle of
the planet (and in excess relative to the other major volatile elements; Table
1) with respect to the atmosphere. Consequently, at equilibrium, the fO2 of
the atmosphere is equivalent to that in the mantle, given by:

$FeO(l)+\frac{1}{4}O_{2}(g)=FeO_{1.5}(l),$ (4)

[74] exploited this property to determine the speciation and mass of an
atmosphere in equilibrium with a magma ocean of peridotitic composition;
assuming Fe3+O1.5/(Fe3+O1.5 \+ Fe2+O) = 0.037 (that of the present-day Earth’s
mantle), would set the fO2 of the atmosphere in equilibrium with the magma to
$\sim$IW222IW = the $f$O2 set by the iron-wüstite buffer; FeO = Fe + 1/2O2.
[[, see also]]hirschmann2022magma. At 2173 K, the temperature at which
peridotite is fully molten at 1 bar pressure, this corresponds to $\sim$10-6
bar of O2. Its importance lies in fixing the relative fugacities of
atmospheric species that can be related by fO2 alone, such as;

$H_{2}(g)+\frac{1}{2}O_{2}(g)=H_{2}O(g),$ (5a)
$CO(g)+\frac{1}{2}O_{2}(g)=CO_{2}(g),$ (5b)

The CO/CO2 and H2/H2O ratios of the atmosphere, as well as their total
pressures, are key in dictating the physical and chemical stability of
atmospheres around rocky planets [[, e.g.,]]lichtenberg2021vertically. The
H2/H2O ratio, in particular, influences the fugacities of CH4 and NH3, both of
which are instrumental in the production of amino-acids by spark discharge in
the presence of liquid H2O, as investigated in the Miller-Urey experiment
[77]. Their homogeneous gas phase reactions can be written;

$CO(g)+3H_{2}(g)=CH_{4}(g)+H_{2}O(g),$ (6a) $N_{2}(g)+3H_{2}(g)=2NH_{3}(g),$
(6b)

The prevalence of such gases depends on i) the absolute abundances of H, C and
N in the atmosphere and ii) the atmospheric H2/H2O ratio, which is, in turn,
governed by the temperature, pressure and fO2. In order to investigate the
effect of fO2 on the chemistry of atmospheres produced on the early Earth in
equilibrium with a magma ocean, the BSE abundances of the CHONPS elements
(Table 1) are used together with the relative fugacities of stable gas species
given by the equilibrium constants (Chase, 1998) of their homogeneous gas
phase reactions (e.g., eqs. 5a, 5b, 6a, 6b) and their solubilities in silicate
liquids [[, see]and references therein]bower_etal2022. The partial pressures
(fugacities) of the major atmosphere-forming species at 2173 K at three
different oxygen fugacities [[, expressed relative to the iron-wüstite
buffer;]]oneillpownceby1993 in equilibrium with an Earth’s mantle-sized magma
ocean are shown in Fig. 3.

Figure 3: Equilibrium speciation of atmospheres at 2173 K formed around an
Earth-sized planet in equilibrium with a magma ocean of mass equal to that in
Earth’s mantle (4.2 $\times$ 1024 kg) in the system CHONPS with abundances for
the bulk silicate Earth (BSE) as per Table 1. The atmospheres are produced
given solubility laws for CO [80], CO2 [66], H2O [64], H2 [68], N2 [65], SO2
[73] and S2 [81] and their equilibrium gas speciation solved at three oxygen
fugacities relative to the iron-wüstite (IW) buffer, +3, 0 and -3. Total
pressure in the atmosphere is shown below the respective pie chart.

Under reducing conditions ($\Delta$IW-3), an early Earth atmosphere would have
had subequal amounts of CO and H2 [[, cf.]Fig. 3]sossi2020redox, while CH4 and
NH3 are present at 0.35 bar and 0.013 bar, respectively. By contrast, S is
essentially absent, with the most abundant S-bearing species, H2S =
2$\times$10-5 bar. The total pressure at the IW buffer defines a minimum,
$\sim$120 bar, as, at lower fO2, CO is joined by H2 and at high fO2, CO is
replaced by CO2 and significant quantities of SO2 ($\sim$80 bar, but also
$\sim$1 bar each of SO and S2) are present in the atmosphere, a result
emphasised by [82]. This reflects the low solubility of SO${}_{4}^{2-}$ in
silicate melt at high fO2 [[, cf.]]boulliungwood2023sulfur. Water is not a
major component in any atmospheric type as its high solubility ensures it
remains dissolved in the silicate liquid [[, see
also]]bower_etal2022,sossi_etal2023, with the implication being that the
$\sim$270 bar equivalent mass of H2O present in the oceans today is a product
of the long-term evolution of the Earth (see section 3), starting with the
solidification of the magma ocean [78, 83, 84]. That is, owing to the lower
solubility of H2O in solid phases relative to silicate liquid, crystallisation
tends to expel water from the magma ocean into the atmosphere. However, the
degree to which H2O degasses depends upon the mode of crystallisation, where
convective lock-up of the magma ocean when the melt fraction is $\sim$0.3
prevents further outgassing of water, whereas efficient crystal settling leads
to a surface magma ocean, allows water to escape [78]. The sum of partial
pressures of phosphorus-bearing species are low, ranging from 2$\times$10-3
bar at $\Delta$IW-3 (mainly P, PO2, PO and P2) to 3$\times$10-6 bar at
$\Delta$IW+3 (chiefly PO2). Their low contribution results from the non-
ideality of P2O5 in silicate melts [[, $\gamma$P2O5
$\sim$10-10;]]suito1981phosphorus,fegley2023chemical and the P abundance in
the BSE (87 ppmw; Table 1).

The relative abundances of gaseous species (Fig. 3), change markedly as the
atmospheres cool and components condense and are extracted from the gas phase
[74, 87]. Notably, CH4 and NH3 become increasingly stable down-temperature at
the expense of H2 and CO. However, the magma ocean stage is likely to supply
the initial budgets of elements to the surface that may then become available
to prebiotic synthesis pathways. Greater fractions of S and N are present in
more oxidising atmospheres, whereas P- and H-bearing species are stabilised in
the gas phase at reducing conditions, while C remains volatile throughout the
fO2 range explored. As such, essential prebiotic molecules, such as HCN, can
only form when the partial pressures of H2, CO and N2 are high enough, with
its presence being favoured at reducing conditions, together with CH4 and NH3.
This realisation led to the notion that transient impact events on the early
Earth could have locally produced reducing atmospheres [88, 89, 90, 91].
Nevertheless, whether these gases can contribute to prebiotic reactions hinges
upon the photochemical stability of these species as atmospheres around rocky
planets cool over time.

##  3 The role of the convection regime

###  3.1 Convection regimes On Earth and other planets

Following accretion, rocky planets cool down and crystallize from a magma
ocean, leading to the onset of solid-state mantle convection. Different
interior dynamics give rise to distinct surface expressions, including
habitability. Mantle convection and its surface expressions have a profound
impact in the thermal and compositional evolution of a planet, as mantle
dynamics link various layers of the planet and drive evolution. In particular,
it effectively links the deep interior (down the the planetary core) with the
surface/atmosphere through heat transfer and the effect of the magnetic field
(see section 3.5). Other main expressions of this link include volcanism, and
the resulting outgassing of volatile species into the atmosphere, that in turn
affects atmospheric mass and surface conditions through the greenhouse effect.
In conjunction with other processes (photochemistry, escape, for example),
they contribute to the evolution of planetary surface conditions. Together,
they effectively produce a coupled system with complex feedback processes,
which we are only starting to grasp [[,
i.e.]]schubert1989coupled,lenardic2008climate,driscoll2013,gillmann2014atmosphere,gillmann2022,krissansen2021,ortenzi2020,foley2016.
Here, we focus on the internal component of the system.

Various convective or tectonic regimes have been proposed to exist for
different interior states (see Fig. 4, aligning with distinctive observed
features. Notably, the Earth presently exhibits plate tectonics [98, 99, 100,
101]: namely, its outermost rigid shell, the lithosphere, is fragmented into
several sizable tectonic plates, gradually shifting atop a mechanically weaker
layer called the asthenosphere (part of the upper mantle). Along convergent
boundaries, at subduction zones, cold and dense oceanic plates sink into the
mantle. This subducted material is balanced by the generation of new oceanic
crust along divergent margins and the growth of the lithosphere due to heat
diffusion. These dynamics are intrinsically related with mantle convection,
i.e. the gradual, creeping motion of the Earth’s solid mantle.

Figure 4: Schematic depiction of the primary convection regimes observed or
theorized on terrestrial planets in the Solar System. The three end-members,
in this representation, are Plate Tectonics (PT), Stagnant Lid (SL) and
Plutonic-Squishy Lid (PSL). Episodic Lid (EL) is an intermediate regime
between PT and SL. While the current regimes of Earth, Mercury and Mars are
well defined, Venus’ is still uncertain and could range from EL to PSL, or may
have transitioned recently (a few 100s Myr) into SL. Mercury is in a later
stage of SL compared to Mars because it has cooled down significantly faster.
The state of the convection regimes of the planets should not be considered
static. Instead, it evolves with time, and the figure can be interpreted as a
snapshot of their present-day state. The axis are not fully independent and
should be thought of as possible observables. The convection regime figures
are adapted from [102].

Advances in numerical modeling have significantly enhanced our understanding
of plate tectonics, while also revealing other tectonic regimes that can
explain observations from other Solar System bodies. These alternative regimes
include the stagnant-lid (as observed on Mars, Mercury or the Moon), where a
planet is covered by a single plate [103, 104, 105], the episodic-lid regime,
characterized by a predominantly motionless lithosphere occasionally
undergoing overturning into the mantle [106, 107, 108, 109, 94, 110], the
ridge-only regime, where a large ridge emerges, slightly compressing the
surrounding lithosphere without forming subduction zones [111, 112], and the
plutonic-squishy lid regime, distinguished by a thin lithosphere and several
small, strong plates or blocks separated by warm and weak regions generated by
plutonism [113, 114]. Venus is currently believed to have either an episodic
or a plutonic-squishy lid, yet the exact nature of the tectonic regime on
Venus remains mysterious [22, 115].

On Earth, substantial evidence supports the notion that plate tectonics played
a pivotal -perhaps indispensable- role in the genesis and development of its
atmosphere, oceans, continents, and life [[, e.g.,]]dehant2019geoscience,
sobolev:2011nat, Stern:2016cc, zaffos2017,stern2023. Increasing evidence
indicates that plate tectonics, as a global tectono-magmatic system, has two
significant impacts: (1) it governs the cycling of essential nutrients between
various reservoirs such as the geosphere, oceans, and atmosphere through
diverse tectonic, magmatic, and surface processes [121], and (2) the gradual
movement of tectonic plates, along with associated topographic alterations and
consistent moderate magmatic activity, shaped and fostered biological
evolution [[, e.g.,]]leprieur2016, pellissier2017, Zerkle2018.

It is worth pointing out that there are indications that the inception of life
might have occurred on an Earth devoid of plate tectonics. This proposition
stems from zircon palaeointensity data, which indicates unvarying latitudes
for the samples dating back from about 3.9 billion years ago (Ga) to about 3.3
Ga [124], implying that no large-scale surface motion occurred at the time on
Earth [[, contrast for example with]and references therein]lammer2018. The
earliest indications of life on Earth are older than 3.7 Ga [[,
e.g.,]]rosing99.

Despite its significance on Earth, plate tectonics has not been observed on
any other planet. Over the last few years, there has been a notable increase
in the discovery of rocky exoplanets, particularly the ones known as super-
Earths, which are large terrestrial planets with masses up to ten times that
of Earth, roughly up to twice its diameter. The majority of studies indicate
that these planets should experience a mobile-lid regime or plate tectonics
[63, 127, 128, 129]. Nevertheless, some suggest the contrary [130, 131].
Therefore, it is also worth contemplating whether life can emerge and flourish
on planets with different tectonic regimes.

###  3.2 The interior of terrestrial planets and outgassing

The study of how the interiors of terrestrial planets relate to their
outgassing histories is still in its early stages. Surface conditions and
atmospheric composition are controlled by outgassing (together with volatile
sinks like atmospheric escape and weathering), while volcanism and tectonics
share a profound connection. The case of Mars illustrates that stagnant lid
planets could outgas sufficient volatiles to affect the evolution of their
surface conditions [132, 133, 134, 135, 136, 137]. However, it is often
suggested that, on average, volcanism on a planet with plate tectonics would
lead to more pronounced outgassing than on a stagnant lid (albeit not
necessarily more melt production, see [109]). On the other hand, an episodic
lid regime [94] could produce a step-like evolution of surface conditions,
with changes occurring in short, sudden bursts, which may be detrimental to
life. At an extreme level, some volcanic events, such as Large Igneous
Provinces have been suggested to alter surface conditions so strongly that
they could cause the end of planetary habitability [[, see]for
Venus]way2022large.

The composition and release of gases during partial melting and volcanic
activity depend on the mantle’s composition and redox state (section 2).
However, uncertainties exist regarding these factors, even for Earth [139,
140, 141, 142, 143, 144, 74]. The redox state influences the speciation of
volatile elements in magmas, and, given that these species have different
solubilities, the degree to which they are released from magmas (cf. 2, eqs.
5a, 5a).

The redox state of the mantle (fO2) is likely inherited from the composition
of planetary building blocks, and the conditions under which core-mantle
equilibrium took place [[, e.g.]]armstrong_etal2019,deng2020magma. However,
these initial conditions may be affected by the subsequent tectonic regime
experienced by a planet. Notably, it has been proposed that subduction,
through the burial of oxidised components such as H2O, CO2 and sulfates, has
lead to an increase in the oxidation state of Earth’s mantle over time,
although this remains a topic of debate [[,
e.g.,]]Aulbach2016,stolper2019,nicklas2018redox, while other processes, such
as disproportionation of Fe2+ into Fe0 and Fe3+ are also plausible [149, 145].

The pressure at which degassing occurs also plays a significant role in
determining the composition and mass of the released species [150]. In the
case of Venus, with a surface pressure $\sim$90 times greater than Earth’s,
volcanoes are expected to emit carbon-dominated gases [22], whereas on Earth,
volcanoes release more water-rich gases. For large exoplanets with extremely
high atmospheric pressures, fewer volatiles would be released [150, 151].

The size of a planet also also appears as a major factor influencing volatile
history. For instance, escape processes depend in a large part on the
gravitational pull of the planet, with small planets experiencing stronger
loss over time than larger ones [125, for example,]. One must also consider
that different species escape at different rates. Lighter species tend to
escape more readily; for instance Hydrogen. Heavier species tend to be
depleted to a smaller extent. However, energy input from the Sun, atmosphere
structure and composition, and photochemical reactions involving each species
are critical to quantifying accurately the escape rates of each atmospheric
component [152, 153]. This affects volatile retention and the whole volatile
inventory, from accretion to the present-day. The specifics of the escape
processes vary from planet to planet, not only with size but also atmosphere
composition and structure, as well as the presence or absence of a magnetic
field (see section 3.5). This goes well beyond the scope of this work, and the
interested reader is welcome to peruse further review papers, for instance:
[152, 153], [22] (for a look at the interactions leading to the general
volatile inventory) or [83] (for the primitive evolution).

Focusing on internal processes, larger planets also generate more melt over
longer periods. Because the heat of formation scales with volume and the heat
loss scales with area, larger planets cool more gradually compared to small
planets. The composition of the mantle and crust also affects the thermal
evolution. For example, stars with C/O$>$0.65 [[, for about 30% of stars in
the GALAH database,]]spaargaren2023 are expected to be hosting carbon-rich
rocky exoplanets [155, 156] with a graphite lid/crust or a diamond-silicate
mantle. The presence of graphite/diamond would speed up planetary cooling by
10–100 times, leading to convection modes that are sluggish or stagnant with a
low degree of partial melting [[]]unterborn2014,hakim2019. This difference
arises from the 10–100 times higher thermal conductivities of graphite and
diamond than that of common silicates. The final regime of planets that have
lost most of their heat is an inactive stagnant lid (the Moon, Mercury),
although not every stagnant-lid planet is geologically ”dead”. The type of
convection or tectonic regime they undergo is also crucial [159, 114]. Recent
studies have suggested that for stagnant-lid planets, melt production and
volatile outgassing are more favorable for planets within the range of 2–4
Earth masses [160, 161, 96].

###  3.3 Observables

Remaining unknowns regarding the interplay between degassing and tectonic
regimes still impede efforts to reliably link specific atmospheres types
(compositions, species, mass) to a given regime. Future work should aim at
understanding the types of atmospheres that different tectonic regimes can
build and their potential for detection.

Magmatic heat flow has been analyzed for various tectonic regimes [114]. Heat
flow is closely linked to erupted materials and, consequently, to outgassing.
Provided sufficient heat sources, the highest magmatic heat flows are observed
among active stagnant-lid planets with volcanism such as Io, which is the most
volcanically active body in our Solar System [162]. Io’s volcanism efficiently
transports approximately 40 times the Earth’s heat flux from its interior to
its surface [163, 164]. On the opposite end of the spectrum, plate tectonics
and plutonic-squishy lid planets experience the lowest magmatic heat flows.
Despite significant differences and uncertainties regarding intrusion to
extrusion ratios, these regimes are expected to have relatively similar
magmatic activity and outgassing histories compared to other tectonic regimes.
Reliable measurements of the heat transport (through surface heat flow) are a
major tool of planetary characterization of internal structure and dynamics
[165, 115, 166], but remain delicate, even on Earth [167].

###  3.4 Surface and interior water inventory

Assessing planetary habitability involves understanding the water inventory
and its evolution over time. The planet’s water retention depends on its
evolution during the magma ocean stage, as some or most of the water is
retained. On Earth, a significant portion of the delivered water was retained,
becoming a fundamental cornerstone in its evolutionary history, which may not
have been the case for Mars or Venus [[, e.g.]]gillmann2020,gillmann2022.
Collisions could also have affected the planetary climate [169, 170, 171].

Tectonic regimes also govern the availability of liquid water on the surface,
as well as the weathering of silicate rocks and the burial of marine carbonate
rocks into the mantle. These latter mechanisms contribute to climate cooling
by removing CO2 from the atmosphere and the planet’s surface, respectively [[,
e.g.,]]west2005, plank2019. In turn, the recycling flux can deliver back
volatiles to the mantle that can later be incorporated again in the magma.
This mechanism could help prevent depletion of the mantle and ensure
continuous outgassing over geological times.

The presence of water has far-reaching effects on mantle rheology, reducing
its effective viscosity by several orders of magnitude [174, 175]. It also
influences melting processes by lowering the mantle solidus temperature [[,
e.g.,]]green2010water. Surface water plays a crucial role as well:

  1. 1.

Hydrated crust facilitates the lubrication of subduction slabs, enabling
continuous one-sided oceanic subduction [177, 178], a key factor and primary
driver of plate tectonics [179].

  2. 2.

On Earth, after glaciations, the delivery of sediments into oceanic trenches
is enhanced, further lubricating subduction and accelerating plate tectonics
[180].

  3. 3.

As a condensable greenhouse gas, water significantly affects surface
conditions, up to to runaway greenhouse (refer to section 4, [181]).

  4. 4.

The extent of Earth’s surface covered by water significantly influences the
planet’s climate [182, 183, 184, 185, 186, 187, 188], which can be linked to
the thermo-convective evolution of Earth [189].

Therefore, changes in surface conditions affect the planet’s interior. For
instance, water (and other gases) influences surface temperature, which, in
turn, affects the convection patterns [93, 94]. The presence of life and
interaction with water further complicates the situation. For instance, it has
been suggested that life may trigger plate tectonics [190]. These complex
interactions across multiple levels pose challenges for modeling realistic
planetary scenarios.

Future research endeavors should therefore also strive to integrate and
comprehend the interconnections and feedback loops among tectonics, water
circulation, landscapes, climate, and life within diverse tectonic regimes. By
doing so, a more complete and detailed depiction of planetary habitability can
emerge.

###  3.5 Magnetic field generation: an illustration of the far-reaching
effects of mantle dynamics

The generation of a magnetic field by a terrestrial planet is a good example
of the links between the interior of a planet and its surface conditions, as
well as an argument in favor of considering the planet as a complete system.
Firstly, magnetic fields are generated in the planetary core by a magnetic
dynamo.Secondly, enough heat needs to be extracted from the core, through the
mantle and out of the solid planet, in order for a dynamo to operate, again
demonstrating the importance of considering a planet as a system. Thirdly, the
presence of a magnetic field affects the interaction between solar radiation
and the planet’s atmosphere. In fact, there have been suggestions that a
magnetic field may be necessary for a planet to be habitable or, at least, for
life to evolve toward a more complex state [1].

The significance of the magnetic field for supporting life is underscored by
several arguments. Firstly, it provides protection to a planet’s surface by
shielding it from harmful solar radiation. Additionally, the magnetic field
acts as a barrier against some atmospheric (especially water) loss [191].
Among the planets in our solar system, Earth is the sole planet with a self-
generated magnetic field. Mars, on the other hand, bears traces of an extinct
magnetic field, while in its ancient crust, which would have been active
during the Noachian era, roughly 4 to 3.8 billion years ago, a period when the
planet may have been less arid than it is today. Venus, in contrast, exhibits
no evidence of an intrinsic magnetic field in recent history, nor are there
any observations indicating remnants of one, which is surprising, given the
usual assumption of similar structure and composition compared to Earth [192],
and may indicate a radically different thermal evolution.

On Earth (and probably many rocky planets), the iron core is initially fully
molten [193] and its cooling is governed by the heat extraction by mantle
dynamics. Producing a dynamo requires planetary rotation and convection in a
conductive layer. Convection can occur in the liquid part of the core,
triggered by the core/mantle temperature contrast or chemical differentiation.
As cooling proceeds, solid iron sinks to form a solid inner core while lighter
elements (such as H, C, O, Si, S, for example) rise. The motion of
electrically conducting iron in an existing magnetic field induces currents
generating a new magnetic field.

In both cases, magnetic fields are an expression of the thermal history and
internal structure/composition of terrestrial planets. The rate of heat
extraction is required to be large enough for convection to take place. On
Earth, it is believed that a minimum heat flow out of the core of 6.9 TW [194]
is required , which is nearly within error of Earth’s estimated CMB heat flow
of 7–17 TW [195]. If Venus had an Earth-like core, the critical heat flow
value would probably be around 5 TW [196]. However, different core properties
are poorly constrained, crucially the thermal conductivity, [192].

In general, stagnant lid regimes are less effective at extracting heat from a
planet’s interior compared to mobile lid regimes. Consequently, a mobile lid
regime, such as plate tectonics, is more likely to sustain long-lasting
convection in the core. This disparity may be exemplified by the contrasting
situations of Earth and Mars [116], where Mars has not generated a magnetic
field for approximately 4 billion years, due to its still liquid core [197]
and the lack of efficient heat extraction after the primitive period. On
Venus, other reasons have been proposed to be behind the absence of a magnetic
field. These include an already fully solidified core [198] or a stably
stratified core [192, 199] resistant to convective motion (possibly due to the
absence of remixing caused by a giant impact, such as the Earth is thought to
have had leading to the Moon formation).

However, recent observation of the atmospheric escape rates of O+ ion on Mars,
Venus and the Earth suggest that they are rather similar. Loss rates ranges
from 2-4 $\times$ 1024 O+ s-1 for Mars, to 1-4 $\times$ 1024 O+ s-1 for Venus
and 5 $\times$ 1024 O+ s-1 to 5 $\times$ 1025 O+ s-1 for Earth [[, see]and
references therin]ramstad2021. This implies that perhaps the magnetic field is
less important than previously thought regarding atmospheric escape.

It has been suggested that a magnetic field can extend much further away from
the surface of the planet than the atmosphere itself, which causes a
magnetized planet to intercept a larger portion of stellar energy when
compared to an unmagnetized one [201]. That additional energy would power
inflated escape rates, possibly higher than for unmagnetized planets [202,
e.g.]. This postulate has been challenged by [191], who point out poorly
constrained return fluxes and the variability of escape rates with the solar
flux and geodynamo even for a single planet [200].

An undeniable effect of self-generated magnetic fields is that they modify the
preferential escape location: ions flow along magnetic field lines under the
effect of the solar EUV flux. With an induced magnetic field, ions could be
lost in the magnetotail, as magnetic lines are open. With self-generated
magnetic field, magnetic limes in the tail reconnect to the planet, preventing
ions escape. However, the cusps of the magnetic field contain open field
lines, and allow solar radiation to excite ions deep in the atmosphere [203]
and accelerate them downtail through the lobes of the field, leading to their
escape. This was further illustrated by models of magnetized planets with a
high obliquity.

It is possible that what differs between magnetized and unmagnetized planets
is the response of magnetospheres/ionospheres to intense events, like Coronal
Mass Ejections (CME), which are far more numerous during planetary early
evolution [204]. To summarize, current research casts doubt onto the classical
view of the magnetic field as a shield against escape, but highlights the
complexity of its interactions with solar radiation. The net result for
terrestrial planets remains uncertain.

##  4 The cessation of habitability

###  4.1 Liquid water and carbon cycling

The breakdown in one of the many processes that enable long-term habitable
conditions on Earth can drive the cessation of habitability. Liquid water, an
energy source, CHONPS, transition metals and relevant geochemical conditions
enable life to originate and evolve. The sustenance of a core-dynamo-driven
magnetic field protects the Earth’s biosphere from cosmic rays and charged
solar particles. The steady interior evolution and plate tectonics drive
outgassing and recycling of volatiles on Earth. The presence of oceans, the
hydrologic cycle and the carbonate-silicate cycle (inorganic carbon cycle)
allow for temperate climates by regulating CO2, a vital greenhouse gas. Among
these factors, the presence of liquid water on the surface and the geochemical
cycling of CO2 are key in ensuring habitable conditions on billion-year
timescales.

Unlike Venus, where most of the near-surface CO2 is present in the atmosphere,
on present-day Earth, more than 99% of near-surface CO2 is expected to be
locked up in the form of carbonates, courtesy of the carbonate-silicate cycle
[[, e.g.
]]urey1952,DonahuePollack1983,wedepohl1995composition,Lecuyer2000,hartmann2012geochemical.
A number of steps are involved in recycling CO2 between the atmosphere and the
interior of Earth [210, 211, 212]. The continental weathering of silicate
rocks [213, 214] and subsequent seafloor precipitation of carbonates locks up
CO2 in solids [215]. Eventual subduction of carbonates removes CO2 from the
atmosphere-ocean system [211]. In the upper mantle, due to metamorphism,
carbonates break down into silicates and release CO2 back into the atmosphere
[216].

To sustain CO2 cycling, several factors are expected to be crucial, including
the hydrologic cycle, fresh silicate rocks, and a planetary orbit within the
habitable zone [217]. The incident stellar radiation (instellation) on Earth
has gradually increased over the past 4.5 billion years because of the Sun’s
evolution on the main sequence of the Hertzprung-Russell diagram. The faint
young Sun paradox suggests that early Earth must have been frozen to maintain
an energy balance between solar irradiation and outgoing longwave radiation
for an early Earth with similar albedo and atmospheric greenhouse effects as
modern Earth [218]. The contrary evidence of an active hydrologic cycle during
the Archean calls for a mechanism to ensure temperate climates during the
Earth’s history [219]. The carbonate-silicate cycle-driven steady decline in
CO2 partial pressure from approximately 0.1 bar (uncertainty range: 0.01–1
bar) during the Archean to a pre-industrial value of 0.28 mbar has been
suggested to be a solution for the faint young Sun paradox [213, 219]. During
the Phanerozoic, the biosphere has further enhanced the efficiency of CO2
cycling by adding an organic component to the carbonate-silicate cycle [220,
221]. Which factors can break down the carbonate-silicate cycle? This question
is key because it invokes a number of processes of habitability cessation that
are directly or indirectly involved in the carbonate-silicate cycle.

###  4.2 Silicate weathering and carbonate precipitation

If silicate weathering were not operational on Earth, the partial pressure of
atmospheric CO2 accumulated due to volcanism would already become 1 bar in 20
Myr [213]. This would rapidly increase the surface temperature due to the
greenhouse effect of CO2. Fortunately, that is not the case on Earth. The
balance between weathering and outgassing on a shorter timescale of 0.1–1 Myr
ensures a quasi-steady state in atmospheric CO2 [222]. On a longer timescale
of 0.1–1 Gyr, all planetary CO2 reservoirs can attain equilibrium and prevent
a runaway greenhouse state [211, 223]. The negative feedback of silicate
weathering ensures that the accumulation of CO2 due to outgassing will be
balanced by the loss of CO2 due to silicate weathering, thereby maintaining
moderate amounts of steady-state CO2 on shorter timescales. On longer
timescales, the intensity of silicate weathering must gradually increase with
stellar luminosity.

The intensity of weathering is limited by the availability of weathering
reactants, i.e., water, silicate rocks and CO2 [213]. This leads to three
limits of weathering [214, 172, 224]. If fresh silicate rocks cannot be
supplied by plate tectonics, weathering becomes rock supply-limited $W_{\rm
su}$ [172, 223]. However, plate tectonics may not be necessary for carbonate
burial and recycling [[, e.g., CO2 cycling for stagnant-lid
planets]]foley2018,hoening2019. From the perspective of the limited transport
of hydrologically important flowpaths to fresh minerals leading to dilute
water streams, supply-limited weathering is also called transport-limited
[214, 227]. Weathering becomes kinetically limited $W_{\rm ki}$ when the time
available for weathering reactions is shorter than the reaction equilibrium
timescale [213, 214]. If there is limited availability of water (e.g., arid
climates), weathering becomes thermodynamically limited $W_{\rm th}$ [224,
228, 229]. Contrary to the kinetic limit, at the thermodynamic limit,
weathering decreases as a function of temperature, resulting in a positive
weathering feedback, whose impact on climate stability is not understood yet
[229]. To accurately model CO2 cycling of temperate planets beyond modern-
Earth conditions, [230, 229] demonstrate the need to use a multi-regime
weathering equation [[]adapted formulation]maher2014:

$W_{\rm total}=\left(\frac{1}{W_{\rm th}}+\frac{1}{(\frac{1}{W_{\rm
ki}}+\frac{1}{W_{\rm su}})^{-1}}\right)^{-1}.$ (7)

With such a formulation, the [213] kinetic weathering regime is recovered when
$W_{\rm ki}<W_{\rm su}$ and $W_{\rm ki}<W_{\rm th}$, which is largely the case
for modern Earth. The weathering is rock-supply limited when $W_{\rm
su}<W_{\rm ki}$ and $W_{\rm su}<W_{\rm th}$, and the weathering is at the
thermodynamic limit when $W_{\rm th}<(1/W_{\rm ki}+1/W_{\rm su})^{-1}$.

The type of silicate rocks also affects the intensity of weathering. For
example, ultramafic and mafic rocks contribute up to 10 times more to
weathering than felsic rocks [231, 229]. Rocky exoplanets are expected to show
a large diversity in silicate rock types, as studies using stellar elemental
abundances suggest [232, 154]. Other planets may feature a wide range of
gas/surface chemical reactions, that may or may not be conducive to
habitability. For example, it is thought that oxidation reactions on Venus
could consume or buffer oxidized components of the atmosphere, such as H2O,
SO2 or CO2 [233, 22].

As the products of silicate weathering reactions are transferred to oceans by
rivers, carbonates precipitate and settle on the seafloor for eventual
subduction into the mantle. Carbonate precipitation is the following step
after silicate weathering. On present-day Earth, the precipitation of calcium
carbonates (calcite, aragonite) occurs almost instantaneously compared to
silicate weathering [211]. The precipitation of carbonates is generally
neglected in carbon cycle box models because it is not the rate-limiting state
among the two. However, even when continental silicate weathering takes place
but the oceans are too deep, carbonate precipitation cannot occur. Below a
certain ocean depth, known as the carbonate compensation depth (CCD),
carbonates are unstable due to high pressures and dissolve back into the ocean
[234, 215]. To ensure the availability of carbonates for subduction, in
principle, the CCD should be deeper than the ocean depth, which is the case
for modern Earth. The CCDs for the carbonates of calcium, magnesium and iron
become deeper at a higher $P_{\rm CO_{2}}$ and a higher temperature, thereby
favouring CO2 cycling for diverse carbonates [235]. In the absence of
continents, seafloor weathering and CO2 dissolution on ocean worlds can also
stabilise the long-term climate [236, 227, 237]. The total liquid water
inventory and topography of terrestrial planets (which depend on outgassing
and mantle dynamics) can strongly affect the drawdown of CO2.

Figure 5: CO2 volume mixing ratio (VMR${}_{\rm CO_{2}}$) of two exoplanet
types as a function of instellation normalised to that of modern Earth
($S_{\rm eff}$ =1). A wet planet, i.e., an Earth-like planet with surface
water oceans, carbonate-silicate cycling and an N2-dominated atmosphere can
potentially maintain a constant surface temperature $T_{\rm surf}=288$ K by
decreasing VMR${}_{\rm CO_{2}}$ in response to an increase in $S_{\rm eff}$
[[, calculated using the climate model of]]kadoya2019. The inner HZ edge is
shown as an extended zone because 1D and 3D climate models predict the inner
edge to be between $S_{\rm eff}$ = 1.05–1.2 [15, 239, 240, 241, 242]. At
higher installations beyond the inner HZ edge, water evaporation/loss followed
by near-surface CO2 desiccation will make the atmosphere of the wet planet
CO2-dominated exhibit a jump in VMR${}_{\rm CO_{2}}$ [243, 244, 230]. For a
CO2-dominated dry planet, the inner edge of HZ is not expected to affect
VMR${}_{\rm CO_{2}}$.

###  4.3 Observational tests

Inside the habitable zone (HZ) for a wet planet, the carbonate-silicate cycle
can be assumed to result in a constant surface temperature (e.g., $T_{\rm
surf}=288$ K, Figure 5). Such a stable temperature can be sustained by
ensuring a higher $P_{\rm CO_{2}}$ at lower instellation due to a lower
intensity of silicate weathering and a lower $P_{\rm CO_{2}}$ at higher
instellation due to a higher intensity of silicate weathering. At sufficiently
high instellations [[, higher than 1.05–1.2 $S_{\rm eff}$, where $S_{\rm
eff}=1$ denotes modern Earth
insolation,]]kopparapu2013,yang2014,wolf2014,wolf2015,way2018, the weathering
feedback cannot mitigate climate warming (above a fixed energy deposition
limit). This leads to the evaporation of water and an increase in the
atmospheric water vapour abundance beyond the regular water vapour positive
feedback [217]. The intricate balance between climate feedbacks and
geochemical cycling is key to avoiding such an irreversible increase in
surface temperature by runaway greenhouse due to the water vapour feedback
mechanism, or by an additional greenhouse gas that pushes the surface
temperature beyond the critical point of water on its own [[, the Simpson-
Nakajima and Komabayasi-Ingersoll limits,
respectively]]simpson1929,komabayasi1967,ingersoll1969,nakajima1992. The
orbital distance at which the runaway greenhouse occurs is thought to define
the inner HZ edge [16, 15]. A bimodal distribution in the CO2 volume mixing
ratio (VMR${}_{\rm CO_{2}}$) across the inner HZ edge for wet rocky planets is
statistically observable [243, 230, 249]. [243] suggested that 20 planets
could be sufficient to identify the VMR${}_{\rm CO_{2}}$ jump at the inner HZ
edge (Figure 5). On the other hand, [250] made statistical predictions for the
planets inside the HZ. If the probability of observed planets matches a log-
uniform distribution of $P_{\rm CO_{2}}$ regulated by the carbonate-silicate
cycle, future telescopes need to observe at least 83 HZ planets [250]. [251]
suggest that a high-precision transit photometry survey, such as ESA PLATO
[252], can observationally identify the inner HZ edge VMR${}_{\rm CO_{2}}$
jump for a sample of 100 exoplanets with at least 10% exhibiting runaway
climates.

It is important to note that an alternative climate state with CO2 as
clathrate hydrate or liquid CO2 can also be stable for billions of years
[253]. The critical absorbed flux can be crossed due to an increase in stellar
radiation or changes in albedo, for instance. The moist atmosphere has a limit
to how much energy it can radiate into space, which will further increase the
surface temperature until the atmosphere radiates in the near-infrared. It is
possible that a runaway greenhouse could lead to partial melt of the surface
for large enough surface liquid water inventory [[, about 1400 K for a few
tens of bar,]]zahnle2007. This positive feedback can persist with almost no
bounds leading to a runaway greenhouse. Venus is suspected to have undergone
such a runaway greenhouse during its past [206, 181]. Even on Earth, such a
runaway greenhouse is expected to occur in the future as the Sun becomes
brighter on its main sequence evolution.

##  5 Current and future missions

###  5.1 Solar System missions

To this date, no space mission has ever found definitive evidence for life in
the solar system, beyond Earth, but habitable environments have been suggested
for specific locations and times. Ancient Mars, ancient Venus, and outer solar
system icy satellites are all interesting targets. Direct observation of
habitable conditions is difficult, and space missions tend to investigate
concomitant properties of the planet and environment, to help understand
planetary evolution.

  * •

Mars is the most explored planet in the Solar System, apart from Earth. After
negative results of the Viking mission [255], it was recognized that present-
day Mars is not favorable to life, despite minute signs that habitability
might not be far off martian conditions. Any habitable phase would have
occurred more than 3.8 Gyr ago.

Further exploration operated under the ”follow the water” guideline and
attempted to characterize the ancient geology of Mars and put time constraints
on water-related, geological and mineralogical features [[,
e.g.]]carr2007,bibring2006,squyres2005. Mars, once, had more water (liquid, at
times) than at present, even if it remains uncertain whether habitable
conditions were stable or episodic, or if the climate was warm or cold [[,
see]for instance]wordsworth2016,bishop2018.

Further investigation has targeted interactions between the interior and the
atmosphere. Phoenix (which arrived on Mars in 2008) has investigated polar
region rhegolith and found that it has interacted little with liquid water in
recent history [261]. The Curiosity rover (landed in 2012) and Perseverance
(landed in 2018) looked into the possible subsurface origin of methane [262,
263], and provided data regarding surface mineralogy, past traces of water,
the lower atmosphere and geology.Finally, Tianwen-1 (China National Space
Administration, CNSA) landed in 2021 and notably explored the relation between
surface features, the past presence of water and the nature of the martian
surface.

Insight (2018-2022) aimed to cover the deep interior of Mars and its
structure, composition and dynamics, in particular using a seismometer [264,
197]. The mission recorded hundreds of Mars-quakes, and constrained the size
of the core and its fraction of light elements.

Future Mars-related astrobiology missions include multiple plans for a Mars
Sample Return (NASA), as many questions require direct analyses of martian
material [265, 266]. Plans for the Tianwen-3 mission (CNSA) also include
sample return (launch scheduled for 2028). Finally, the two-part Exomars
mission (2016 and 2028 launches) attempts to investigate biosignatures, water
distribution, carbon isotopes and their possible biologic origins.

  * •

Venus has been relatively little explored compared to other targets in the
solar system.

Early remote observation, Mariner 2 and the soviet Vega and Venera landers (in
the 60’s and the 70’s) revealed the crushing 92 bar surface pressure, and
$\approx$740 K surface temperature caused by its massive CO2 atmosphere [[,
see]for a detailed overview]o2022venus.

Ancient habitable Venus has nonetheless been suggested [268], but is debated
[[, see]]gillmann2022,westall2023 and depends on very specific atmospheric
conditions [269]. Present-day habitability of the atmosphere near the cloud
layer has been suggested [270] but remains unconfirmed, while recent claims of
the observation of phosphine [271] in the atmosphere have again triggered
heated debate [272, 273, 274, 275, 276]. Beyond invaluable but imprecise in-
situ measurements (Vega, Venera), Magellan (1990) and Venus Express (2006)
have been the source of most of the current scientific knowledge of the
surface and atmosphere of Venus [267]. Additional observation was performed by
Akatsuki (2015) and the Parker Solar Probe (2018), as well as other missions
en-route toward other targets [277].

There is a general consensus that we still know very little about Venus’
interior structure and composition, and lower atmosphere. Yet, the evolution
of Venus’ interior and how it may have caused a divergent evolution compared
to Earth are critical research avenues for habitability.

This caused a resurgence of interest in Venus with several high-profile
missions planned in the next decades: DaVinci and VERITAS (US), EnVision (EU),
VENERA-D (RU), Shukrayaan-1 (India), VOICE (China).

All six missions are designed to investigate the past of the planet and the
processes that affect the evolution of its surface conditions and volcanic
processes. Ancient traces of water are also investigated, throughout the
atmosphere, and near possibly ancient material [[, the
tesserae]]byrne2021venus,whitten2021venus,smrekar2018venus. The main goal of
those missions is not just to know if Venus was habitable, but to understand
why it is not at present-day, so that this knowledge can be applied more
broadly [280, 22].

  * •

The icy moons of Jupiter and Saturn are considered important targets for
astrobiology missions. The Cassini Huygens mission opened the door to further
exploration of the satellites of giant outer solar system planets. The landing
on Titan and discovery of its surface and liquid hydrocarbon lakes has seen
suggestion of methanogenic life [281]. Io, Enceladus, Ganymede and Europa are
all suspected of harboring (possibly habitable) interior liquid water oceans
beneath their icy surface [282, 283, 284, 285]. Future missions include ESA’s
Juice toward Europa, Ganymede and Calisto [286]. Europa is also the target of
NASA’s Europa Clipper [287]. NASA’s second mission, Dragonfly [288, 289] would
place a rotorcraft on Titan to study biosignatures and chemistry.

###  5.2 Exoplanet missions

In the context of detecting and, even more importantly, characterizing the
(atmospheric) properties of temperate, terrestrial exoplanets it is important
to consider a few aspects:

  * •

M-type stars vs. Solar-type stars: Depending on the effective temperature and
luminosity of the host star333Note that we are primarily concerned with main-
sequence stars here., the orbital period (and hence mean orbital separation)
of habitable zone planets varies significantly, ranging from up to a few
hundred days for stars similar or slightly hotter than the Sun, to less than a
few tens of days for cool, low-luminosity M-type dwarf stars. This has
important implications, which techniques can be used for atmospheric
investigations of the objects.

  * •

Transit spectroscopy and secondary eclipse observations vs. direct detection
methods: Currently, the most successful approach to observe the atmospheres of
exoplanets are transit spectroscopy and secondary eclipse measurements [[, for
an introduction, see,]]seager2010. These techniques have a strong bias towards
large, close-in planets. And while the James Webb Space Telescope (JWST) and
also ESA’s Ariel mission [[]]tinetti2018 will revolutionize our understanding
of exoplanet atmospheres for hot/warm gas dominated planets, they will not be
able to probe a large sample of rocky, temperate exoplanets. In the most
favorable cases, which are small planets transiting very low-mass stars [[,
such as the well-known Trappist-1 system,]]gillon2017, JWST may be able to
tell if these objects have an atmosphere at all, which will be an important
milestone in exoplanet science [[, for first results,
see,]]greene2023,zieba2023, but an in-depth characterization and comprehensive
inventory of atmospheric constituents will likely be out of scope [[,
e.g.,]]krissansen-totton2018,lustig-yaeger2019. The reason is very simple:
despite the large 6.5-m primary mirror, the favorable planet-star brightness
ratio, the proximity of the Trappist-1 system of only $\approx$12 pc, and the
ability to observe and stack several transits due to the very short orbital
periods of the planets, the signal-to-noise of the data will remain limited.
This is the reason why, in the future, a main focus will be on direct
detection methods that are largely independent from orbital properties and do
not require an exoplanet to transit its host star444For completeness we
mention here the ambitious and inspiring _Nautilus_ project that seeks to
search for biosignatures in the atmospheres of hundreds of transiting
terrestrial exoplanets with a fleet of low-cost spacecraft [[]]apai2019AJ.;
the vast majority of exoplanets does – in fact – not transit in front of their
stars!

  * •

Reflected light vs. Thermal emission: Accepting that direct detection
techniques are a promising way to probe many of the temperate, terrestrial
exoplanets within 10-25 pc from the Sun, one needs to realize that there are
two ways to directly detect light from these objects: one can either probe the
starlight reflected by the planets (this is typically done at UV, optical, and
near-infrared wavelengths between $\sim$0.2 and $\sim$2.5$\;\mu$m) or one can
probe the planets’ intrinsic thermal emission (which, for temperate objects
similar to the terrestrial planets in the Solar System, can be done at mid-
infrared wavelengths between $\sim$3$\;\mu$m and $\sim$20$\;\mu$m).

With currently available telescopes (on ground and in space), we cannot
directly detect small, habitable zone exoplanets. It’s technically not
feasible as none of the existing facilities or instruments provides the full
combination of required spatial resolution, contrast performance, and
sensitivity. For the required contrast, we remind the reader that in reflected
light at optical wavelengths the Earth is a factor of $\approx 10^{10}$
fainter than the Sun, and at a wavelengths of $\sim$11$\;\mu$m, where the
thermal emission of Earth peaks, the flux difference still amounts to $\approx
10^{7}$ [[, e.g.,]]desmarais2002. For planets orbiting low-luminosity M-dwarfs
these values are typically a factor of 100 lower.

Considering the points listed above, it becomes clear that a number of new
instruments and missions will be needed for a comprehensive investigation of a
large number of temperate terrestrial exoplanets orbiting various kinds of
host stars and across a large wavelength range:

  * •

Probing nearby exoplanets orbiting low-luminosity M-dwarfs in reflected light
will be the primary discovery space of optimized instruments installed at ELTs
[[, e.g.,]]kasper2021. With their 30-40 m primary mirrors they will provide
sufficient spatial resolution to spatially separate the signal from the planet
from that of its host star on the detector, and the combination of extreme
adaptive optics systems with high-resolution spectrographs will deliver the
required contrast performance [[]]snellen2015.

  * •

Probing nearby exoplanets orbiting Sun-like stars in reflected light will
required dedicated space missions. The stringent contrast requirement (see
above) can likely not be achieved by ground-based instruments due to the
disturbing effects from the Earth atmosphere. In fact, a large, single
aperture exoplanet imaging mission was recently recommended in the context of
NASA’s 2020 Astrophysics Decadal
Survey555https://www.nationalacademies.org/our-work/decadal-survey-on-
astronomy-and-astrophysics-2020-astro2020. The main science driver for this
so-called _Habitable World Observatory (HWO)_ is to detect and investigate
$\approx$25 Earth-like planets around nearby Sun-like stars in reflected
light. While the exact wavelength coverage is still under investigation – in
particular the short- and long-wavelength cut-off in the UV and near-infrared
will be carefully examined – the aperture size shall be $\approx$6 m and hence
in between those of the _LUVOIR_ and _HabEx_ mission concepts, which have been
studied in preparation for the Decadal Survey [[]]luvoir2019,habex2020. We
remind the reader that the _Hubble Space Telescope (HST)_ has a primary mirror
of 2.4 m. The oxygen A-band (759–770 nm) is the main diagnostic for _HWO_ to
search for indications of biological activity in exoplanet atmospheres, but
absorption bands of water, and potentially also carbon dioxide and methane,
shall also be accessible. First studies quantifying how well Earth-like
planets can be characterized by a future reflected light mission like _HWO_
have been carried out [[]]feng2018AJ.

  * •

Finally, probing the thermal emission of nearby exoplanets orbiting any type
of host star, requires a large, space-based nulling interferometer. Even with
the 39 m ELT of the European Southern Observatory the thermal background noise
at mid-infrared wavelength for ground-based observations will be prohibitively
high so that only very few of the nearest stars can be imaged directly in the
search of temperate terrestrial exoplanets [[]]quanz2015,bowens2021. From
space, however, a mission like the _Large Interferometer For Exoplanets
(LIFE)_ [[]]quanz2022a could detect hundreds of nearby exoplanets including
dozens similar to Venus and Earth in terms of size and energy influx
[[]]quanz2022b,kammerer2022. An exoplanet’s thermal emission allows for a
robust characterization of its atmospheric properties such as pressure-
temperature profile and composition [[]]line2019 as it is less affected by
disturbing effects such as clouds [[]]kitzmann2011 compared to observations in
reflected light. In the case of a _LIFE-_ like mission, habitable conditions
on an Earth-like planet can readily be inferred from moderate resolution mid-
infrared spectra throughout most of Earth’s history [[]]konrad2022,alei2022
and the mid-infrared wavelength range provides access to a variety of
atmospheric biosignatures gases [[, e.g.,]]schwieterman2018. We emphasize that
the direct detection of temperate terrestrial exoplanets at mid-infrared
wavelengths with the goal to identify habitable – and potentially even
inhabited – exoplanets was given very high scientific priority in the _Voyage
2050_ program of the European Space Agency (ESA) and has been identified as a
potential topic for a future L-class mission in ESA’s science
program666https://www.cosmos.esa.int/web/voyage-2050.

Looking at the list above one can be hopeful that in the coming 20-30 years we
will indeed have a suite of instruments and missions that strongly complement
each other in the search for habitable worlds beyond the Solar System. These
projects are _not_ in competition with each other; in addition to the unique
scientific diagnostic that each of them possesses, the combined power of
reflected light and thermal emission data will provide a truly holistic
picture of temperate terrestrial exoplanets orbiting a wide range of stellar
types in our cosmic neighborhood.

##  6 Conclusions and summary

Habitability is a complex notion dependent on many interwoven processes. The
qualitative definition of habitability being a set of conditions that are
conducive to life is generally insufficient for any precise assessment or
prediction. On the other hand, more quantitative definitions are strongly
biased by the only example of a unquestionably habitable world we possess.

At its core, the habitability of a planet is generally restricted to assessing
the suitability of the conditions at its surface for self-sustaining biotic
reactions to occur; that is, the response of a planetary atmosphere to the
energy input (such as stellar radiation, for example). However, that response
depends on the planet’s time-integrated atmospheric structure and composition,
which are governed by i) the initial availability of life-essential elements
in the planetary body and ii) a myriad of interacting mechanisms and feedback
processes. They can operate both inside the atmosphere (chemistry, dynamics)
or outside (volatile or energy exchange with the planetary interior). The fate
of water is especially important for life (as we know or imagine it, at
least), and constitutes the first order guide for habitability models and
detection.

The relative importance of these different processes changes depending on the
stage of planetary evolution and time. Two of the main drivers of changes are
the evolution of the host star and that of the silicate planet. In this
review, we have discussed some of the most high profile questions regarding
the role of the interior of terrestrial planets in controlling habitability,
such as the conditions for its onset, the role of interior dynamics, and how
it is maintained and could cease.

The nature of terrestrial planets, by definition, is distinct from those of
gas giants that have nearly their full complement of stellar gases. By
contrast, terrestrial planets are relatively ‘dry’ [47]. Therefore, the masses
of atmospheres around terrestrial planets are products of the bulk composition
of the planet and the prevailing pressure-temperature conditions under which
the atmosphere formed and evolved through time. Given that $>$99.95 % of the
bulk silicate mass of planets resides in rocks [38], the proportions of O to
other rock-forming elements, namely, Fe, Mg and Si, control the nature of
atmospheres around rocky planets through the oxygen fugacity, fO2.

As such, a planet’s interior governs the long-term evolution of the atmosphere
by exchanging mass (in the form of volatile species) and heat with it. The net
flux of outgassing (e.g., volcanic exhalation) and ingassing (burial of
volatile elements into the mantle and crust) defines the bulk composition of
the atmosphere (and surface) and thus its water inventory [[,
see]]marty2012,hirschmann2018. We have underlined the importance of the magma
ocean stage in setting the initial conditions that can lead to habitable
surface conditions (section 2). Later, habitability is maintained or fails due
to a delicate balance in volatile exchange. On the Earth, the relevant species
are water and CO2 owing to the relatively oxidised nature of its mantle [[,
$\Delta$IW+3.5 at present in the upper mantle,]]frostmccammon2008, but on
other planets could contain other gaseous species, namely H2 or CO on more
reduced planets, and SO2 on more oxidising planets.

This field of investigation is still in its infancy, as results of the past
few decades have highlighted the vast complexity of the interactions leading
to volatile exchange [[,
e.g.,]]fegleyschaefer2014,hirschmann2018,Zahnle_2020,grewal2020speciation,bower_etal2022.
They involve all layers of the planet, from the core to the upper atmosphere,
in addition to the more obvious atmosphere/mantle interface. This is an
inherently multidisciplinary endeavour that requires a deep understanding of
all the parts of the volatile cycle In short, understanding habitability
requires us to understand planetary evolution as a whole, an equally exciting
and daunting prospect.

Current scientific efforts include many missions especially focused on the
theme of habitability, either in our Solar System or toward exoplanets. It is
tempting to look for the safest targets, selecting objects similar to Earth,
but it is better to cast a wide net and see what comes up. Many missions
attempt this with planets that have very different conditions from Earth, such
as Venus, for example, to understand the mechanisms of planetary evolution, or
for exoplanets, to build-up statistics beyond our small solar system. After
all, there is no guarantee that Earth is the unique blueprint for
habitability. At present, we still understand the topic far too little to make
solid predictions. Instead, we would reap tremendous benefits form observation
of the diversity of possible planetary evolution outcomes.

## Acknowledgements

Funding: Parts of this work has been carried out within the framework of the
National Centre of Competence in Research PlanetS supported by the Swiss
National Science Foundation (SNSF) under grants 51NF40_182901 and
51NF40_205606. PAS thanks the SNSF via an Eccellenza Professorship (203668)
and the Swiss State Secretariat for Education, Research and Innovation (SERI)
under contract number MB22.00033, a SERI-funded ERC Starting Grant ’2ATMO’. KH
acknowledges the FED-tWIN research program STELLA (Prf-2021-022), funded by
the Belgian Science Policy Office (BELSPO). Author contributions: CG directed
this review, organized the paper, wrote sections 1 and 6 and contributed to
sections 2,3 and 5. KH was in charge of section 4. DL was in charge of section
3. SQ was in charge of section 5. PS was in charge of section 2. All authors
discussed, commented and revised the paper. Competing interests: The authors
declare that they have no competing interests.

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