in the thickness of continental crust at the sites of crust generation through
time is inferred from Rb/Sr measurements. The variation of $\delta$18O values
in zircons also marks the shifts in the composition of juvenile crust away
from mafic compositions (Hawkesworth et al., 2020) (adapted from Hawkesworth
et al. (2020)).

As soon as it appeared, maybe multiple times before really gaining a foothold,
early life forms were intimately linked to the volcanic substrates from which
they emerged. Note that all the proposed scenarios for the the emergence of
life referred to above were located in volcanic settings. They were
chemolithotrophic microorganisms obtaining their energy from oxidation of
inorganic materials, such as H2, or through redox alteration of reactive
minerals at the surface of the volcanic rocks (e.g. olivine, pyroxene, metal
sulphides), or chemoorganotrophs oxidising organic substrates, such as dead
organisms or bio-available abiotic organics (Weiss et al., 2016). At this
stage, life was completely dependent upon the mineral substrates. At some
stage, through mutation, life developed the ability to capture energy from
sunlight (phototrophs, (Martin et al., 2018)), thus freeing it of the mineral
substrates and providing it with access to an unlimited source of energy. The
oldest, well-preserved terrestrial rocks dating back to about 3.5 Ga testify
to the already widespread presence of anoxigenic phototrophs in shallow water
environments.

While access to unlimited sunlight gave the phototrophs an energy-rich,
metabolic advantage over the chemotrophs (litho- and organ-) enabling them to
grow faster and create greater biomass, it was the mutation that led to
oxygenic photosynthesis that would change the Earth, both physically,
chemically, and biologically. The timing of this biological invention is
debated with certain geochemical studies suggesting that “whiffs of oxygen”,
i.e. biological oxygen, may have occurred as early as the Paleo-Mesoarchean
($\sim$3.5-3.0 Ga) and during the next 500 Myr (see Fig. 8; Anbar et al.
(2007)). Certainly by 2.9-3.0 Ga there is geochemical evidence for free oxygen
and bioconstructions, such as stromatolites, start becoming common. It has
been suggested that changes in the redox state of the Earth’s surface during
the NeoArchean-Proterozoic transition may have been influenced not only by the
production of oxygen by photosynthesisers, but also by changes in the redox
state of the mantle, that affected the redox of outgassing species from
volcanic eruptions (Wang et al. (2020)), although others have suggested that
there was little change in the oxygen fugacity of the Archean mantle (e.g.
Nicklas et al. (2018)). Whatever the major influences during the Neoarchean,
during this period, huge formations of stromatolites produced by oxygenic
phototrophs occur in Neoarchean formations across the world and, indeed,
dominated the paleontological record for most of the Proterozoic.

Figure 8: Relative concentrations of atmospheric oxygen throughout Earth’s
history, including the classical two-step evolution (orange line, Kump, 2008)
and the more recent model including unstable climate and O2 concentrations
after the first rise of O2 concentrations (Lyons et al., 2014). Possible
“whiffs of oxygen” (blue arrows) during the Neoarchean – and even before
according to Anbar et al. (2007) (adapted from (Lyons et al., 2014))

The turning point was when the oxygen contents in the atmosphere started
slowly to rise between about 2.4–2.3 Ga (the so-called “Great Oxygenation
Event” (GOE), Holland (2002), see Fig. 8). In fact, the gradual accumulation
of oxygen depended upon a delicate interplay of sinks and sources that
probably started as early as about 3.0 Ga (Lyons et al., 2014) and may have
continued until about 800 Ma (Lyons et al., 2021). As noted above, we keep in
mind, however, that there may not have been major changes to the redox state
of the mantle (Nicklas et al., 2018). The production of oxygen was balanced by
its consumption by the initially totally reduced surface of the Earth,
including reduced organic carbon, sulphide minerals, other ferrous minerals,
and reduced volcanic gases, such as H2S, as well as biogenic methane and its
photodissociated product ethane. Concurrent with the gradual oxidation of the
Earth’s surface envelope, including its atmosphere, gradual changes in mantle
buffering resulted in changes in the composition of volcanic gases, more
reduced during the Archean, but more oxidised after the GOE.

While the presence of oxygen in the atmosphere led to elimination of methane
as greenhouse gas, there is still the problem of reducing CO2 in the
Neoarchean atmosphere. The isotopic ratios of carbon in the Neoarchean are
relatively constant, suggesting that there was not a major change in the
burial flux of the total carbon inventory. However, one solution to the
conundrum could be that the flux of buried organic carbon during the Archean
may have been far higher than previously believed (Krissansen-Totton et al.,
2021b). Whatever the processes involved, oxygenation of the atmosphere had the
catastrophic consequence of lowering surface temperatures because the Sun’s
luminosity was still about 20-30% lower than today (as was the case during the
Hadean and Archean, Sagan and Mullen (1972)), which led to the first series of
global glaciations, the Huronian glaciations between 2.45 and 2.22 Ga, when
low latitude continental terranes were covered with glaciers (Bekker and
Holland, 2012). This effectively cut off phototrophic production of oxygen,
increased temperatures due to the accumulation of volcanic CO2 that led to
melting of the glaciers and a sudden massive influx of nutrients that fed the
rapid re-development of the oxygen producing cyanobacteria and the
precipitation of biomediated carbonate deposits, the so-called cap carbonates,
Hoffman et al. (1998). The sudden explosion of biomass production led to a
very large rise in oxygen levels between 2.3 and 2.06 Ga, called the
Lomagundi-Jatuli event, during which oxygen levels almost reached present-day
levels (Melezhik et al., 2013), before falling to very low atmospheric partial
pressure levels of O2 < 0.1% of present atmospheric levels (Lyons et al.,
2014) that reigned for the rest of the Proterozoic until about 0.8 Ga
(Canfield, 1998), as illustrated by Figure 8. It is thought that this long
period of stasis during the Meso-Neoproterozoic was due to the fact that only
the upper part of the oceans were oxygenated while the deep ocean remained
reduced.

However, it was during the Proterozoic that the extra energy available via
oxygenic photosynthesis fuelled the mutations giving rise to eukaryotes, i.e.
organisms with a membrane bound nucleus (prokaryotes, including Bacteria and
Archaea do not have membrane bound nuclei). Indeed, it was only after complete
oxygenation of the entire water column and the stabilisation of nutrient
cycles and environmental conditions could life evolve into the diverse biota
that characterised the terminal stages of the Neoproterozoic (1-0.54 Ga) and
the dawn of the Cambrian period (Planavsky et al., 2015). The Neoproterozoic
period was characterized by extreme biogeochemical and climatic volatility
that led to a series of very severe, low latitude glaciations (Hoffman et al.,
1998; Canfield, 2005; Lyons et al., 2014) resulting in a highly unstable
carbon cycle. This period of multiple “snowball Earths” could have been
triggered or at least influenced by the onset of modern plate tectonics (e.g.
Stern et al., 2018). The global distribution of continents and their positions
relative to the equator likely also had an effect on albedo and therefore
temperatures (Endal and Schatten, 1982; Schrag et al., 2002). At the same
time, changes in oceanic ecosystems led to the rise of eukaryotic algal
primary producers as opposed to the prokaryotic cyanobacteria so prevalent
from the Neoarchean (3.0-2.5 Ga) through the Mesoproterozoic (2.0-1.0 Ga), and
the eventual emergence of metazoan life. Add to this the rise in oxygen
globally and the conditions for significant evolutionary changes to life were
all present (Planavsky et al., 2015).

Today, there appears to be a delicate balance between atmospheric composition,
especially the greenhouse gases CO2 and CH4, the presence of water at Earth’s
surface, subduction of organic matter and carbonates, and subsequent release
of volcanic CO2 back into the Earth’s atmosphere (see Charnay et al. (2020)
and Catling and Zahnle (2020) for recent reviews). It was likely established
gradually throughout the geological evolution of the Earth, possibly as far
back as the Paleoarchean or even before during the Hadean. The exact process
for the recycling of this early carbon cycle is still unknown, but it was most
likely different from modern-day plate tectonics. Initially, the higher CO2
concentration in the Hadean/Archean atmosphere would have offset the lower
luminosity from the Sun (i.e. Kasting, 1988). As solar luminosity increased,
CO2 was gradually removed from the atmosphere by the silicate/carbonate
weathering cycle and sequestration of carbonates and organic carbon into the
crust. In this way, throughout time, climatic stability has been ensured – as
has been the presence of liquid water, one of the prime necessities of life on
Earth.

##  6 Conclusions

The range of possible terrestrial planet evolution is difficult to explore
with only three very particular examples, however, even in our Solar System,
the diversity among the rocky planets is striking. The exact causes for the
clear differences between present-day Mars, Venus and Earth are still
uncertain. It is now suspected that both their early (first few 100 Myr) and
long-term evolution had a role to play. Atmospheric and interior conditions
are the result of billions of years of accumulated and interacting processes
that affect the physical and chemical properties of every planetary layer.

In recent years, our understanding of how magma ocean processes affect the
early evolution of terrestrial planets has grown considerably. Studies show
that in case the terrestrial planets started similarly in a hot molten state,
the iron disproportionation at high pressures in a magma ocean can set a
different initial condition for their subsequent redox evolution. Based on the
assumption that metal-silicate equilibration occurred at the base of the magma
ocean, due to their higher mass, Earth and Venus could have had a deep magma
ocean with a higher abundance of Fe3+ and an oxidizing atmosphere consisting
of CO2, N2 and SO2, while a shallow magma ocean on Mars would be relatively
Fe3+-poor and overlaid by a weakly reducing atmosphere consisting of CO, CO2,
and N2. On all the three planets, the majority of H would be dissolved in the
magma ocean as H2O at an early stage, and then would have outgassed at the
surface as solidification proceeded.

In the case that a steam-dominated atmosphere forms, the subsequent evolution
of the terrestrial planets depends on their orbital distance from the Sun.
Earth and Mars could be sufficiently far from the Sun to keep their water
intact against hydrodynamic escape owing to their short crystallization
period. On the other hand, for Venus, three main different scenarios are
possible: (i) rapid crystallization where water is retained or (ii)
desiccation during a long-lived magma ocean period or (iii) that a high
pressure CO2 dominated atmosphere suppresses water outgassing. In the second
scenario, the Venusian magma ocean can be oxidized by remnant oxygen produced
by photodissociation of water, followed by preferential escape of hydrogen. At
this point, the redox evolution of Earth and Venus may also diverge. If H2O
outgassing is insufficient to form a steam-dominated atmosphere, such an
evolutionary dichotomy would not appear, and early planetary mantles would
contain high abundance of water as discussed in scenario (iii). Planetary
evolution during a magma ocean phase would set the initial conditions in terms
of redox state in the mantle and atmosphere. This can have consequences for
the crystallization history, the total and interior inventory of water and the
climate throughout the entire planetary evolution.

Tectonic regimes can be divided into three main types of convection regimes:
plate tectonics, plutonic-squishy lid, and stagnant lid. Present-day Earth
lies in the plate tectonics regime and features high heat flow and volcanic
production rates. The onset time of modern plate tectonics on Earth is still
under debate and might be related to the presence of life on Earth. The onset
of plate tectonics depends on the rheology of the planetary lithosphere and
the stresses in the convective mantle. Although surface temperature and water
abundance are considered to be important factors affecting the rheology,
whether they favor plate tectonics or not is still debated due to complicated
feedbacks. The presence of subduction zones favors volatile exchanges between
the atmosphere and the mantle. Mars has probably been in the stagnant lid
regime for billion of years with limited geological activity in the last
billion years. The heat flux in the stagnant lid regime is expected to be low,
and this could explain why its core is still in a mostly liquid state, while
not featuring a present-day core dynamo. Its smaller mass and more reduced
interior is expected to limit total outgassing from its interior.

In contrast to Earth and Mars, the tectonic regime of past and present Venus
remains unclear. Present-day Venus does not display signs of Earth-like plate
tectonics, but seems to be more active than Mars and its stagnant lid regime.
Its relatively young surface obscures most signs of prior tectonic and
volcanic activity. A major constraint on Venus’ interior is that at present it
does not feature a global magnetic field. As Venus and Earth are comparable in
size and mass, their differences may originate early on during their
respective accretion and magma ocean stage.

In addition to volatile exchanges between surface and the interior,
atmospheric escape is also a crucial process shaping and driving the evolution
of planetary atmospheres. During the early evolution, hydrodynamic escape and
impact erosion are important processes. Hydrodynamic escape is driven by
strong EUV flux from the young Sun or by heating from a hot surface. Intense
hydrodynamic escape would occur when the upper atmosphere is rich in H-bearing
gaseous species such as H2 and H2O. This process would efficiently remove any
primordial H2-He atmosphere the planetary embryos obtained from the solar
nebula during their formation. Also this process would control the
crystallization rate of a magma ocean located under a runaway greenhouse steam
atmosphere. Impact erosion removes portions of planetary atmospheres. The most
efficient loss occurs on small planets like Mars and/or in the presence of a
surface water ocean. Impacts can also supply material, including volatiles.
The net effect largely depends on the volatile content of the impactors, but
also on the specific impact conditions.

As both impact and EUV flux decrease with time, non-thermal escape processes
gains importance. Non-thermal escape involves photochemistry and charged
particle interactions with the solar wind and among particles. Particles that
obtained sufficient energy from the interaction can escape. Small planets like
Mars are again more vulnerable to loss of heavier volatile species, such as O,
C, N, and noble gases. Water loss, through escape of H and O ions is possible
on all three planets at a comparable, but relatively slow rate. Based on
recent measurements the previous idea that magnetic fields would provide an
efficient shield against non-thermal escape is now being questioned as escape
rates of heavy ions are comparable on Earth (featuring a global magnetic
field), Venus, and Mars (no present-day core dynamo).

Upcoming missions to Venus specifically aim to shed light onto its past
evolution compared to those of Mars and Earth. Likewise, sample return
missions from Mars should lead to a dramatically better understanding of both
Mars interior and surface conditions, as well as whether Mars might be
currently hosting life in the subsurface. The future characterization of a
large number of terrestrial exoplanets will provide further reference points
for the possible evolution paths of rocky planets at all stages of their
histories, including the critical magma ocean phase. The combination
terrestrial planet studies and exoplanet observations will provide both a
detailed and statistic view of terrestrial evolution pathways and will help us
to solve remaining mysteries that we have touched upon within the framework of
this chapter.

## Acknowledgements

The authors thank the anonymous reviewer for his constructive comments that
helped to improve the manuscript. K.H. acknowledges financial support by JSPS
KAKENHI grant No. JP18K13603 and MEXT KAKENHI grant No. JP22H05150. CG
acknowledges that this work has been carried out within the framework of the
NCCR PlanetS supported by the Swiss National Science Foundation under grants
51NF40$\\_182901$ and 51NF40$\\_205606$.

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