284
evolution of life, we have to consider the question
whether and in which way these factors deviated in
pre-Phanerozoic times from the younger Earth's history.
6.5.2 Evolution of the Atmosphere
The Early Atmosphere
Although the atmosphere represents a very small mass
in relation to the total mass of the Earth, it exerts a
dominant influence on the shaping of the landscape,
the existence and evolution of life, weathering of
rocks, transport and deposition of sediments. For this
reason, an atmosphere differing from the present-day
situation must have had a profound impact an all these
aspects.
Ihe evolution ofthe atmosphere was closely related
to the formation of the Earth's crust and hydrosphere.
Most experts hold the opinion that the primordial atmosphere ofthe Earth resembled that of Jupiter. Ihis
planet has sufficient mass to retain light moleeules and
noble gases. In the case of the Earth, however, the
light, volatile components, such as hydrogen and helium, were early lost to space and replaced by an atmosphere mainly caused by outgassing of the Earth's
mantle.
Ihe accretion of planetesimal material, meteoric
impacts, and enhanced radioactive decay raised the
temperature on Earth to such a degree that the primarily more or less homogeneous mass melted. Ihis led to
a differentiation of the accreted mass into core, mantle,
early crust, and a secondary atmosphere derived from
the volatiles of the magma andlor late phases of accretion.
Ihis atmosphere attained a very high density comparable to that of Venus. Its main components were water vapor, carbon dioxide, and nitrogen. Minor constituents included hydrogen, methane, and ammonia.
Light volatiles were continuously lost to space; acids
(HCl, HF) and H 2 S exhaled from volcanoes were
washed out by precipitation. Subsequent rapid heat
loss to space led to cooling of the crust and condensation of water. Evidence for a first hydrosphere on
Earth comes from water-laid sediments ranging back
as far as about 4000 Ma.
The early evolution ofthe Earth, its atmosphere and hydrosphere have been discussed by many workers (e.g.
Schidlowski et al. 1975; Clemmey and Badham 1982; Salop
1983; Holland 1984; Windley 1984; Kempe and Degens
1985; Budyko et al. 1987; Walker andDrever 1988; Wetherill 1990; Condie 1993; Kasting 1993). Further references are
mentioned below.
Chapter 6 Special Depositional Environments
Oxygen in the Early Atmosphere
Whether or not oxygen was present in small quantities
in the early atmosphere is controversial. A limited
amount of oxygen produced by photolysis, i.e., by
dissociation of water vapor as a result of ultra-violet
radiation in the upper atmosphere, was rapidly used up
by oxidation of volcanic gases, ferrous iron dissolved
in sea water (see below), and weathering ofrock-forming minerals. It was not before the plants had started
photosynthesis by using solar energy, water and carbon
dioxide to generate organic compounds and release
molecular oxygen that oxygen could be produced in
larger volumes than were permanently consumed.
However, an excess in free oxygen was only achieved
by the storage of organic marter in sediments. From
then on, the oxygen content of the atmosphere slowly
increased (Fig. 6.18).
Evidence for free oxygen in the atmosphere comes from
paleosols. As long as no oxygen was available, rock surfaces
exposed to the atmosphere lost ferrous iron by leaching.
Later, when the oxygen content ofthe atmosphere had risen,
iron loss became negligible because feme iron cannot be
exported in solution. Of 50 paleosols reported in the literature, 15 could be definitely identified as true palesols (Rhy
and Holland 1998). Paleosols older than 2.2 Ga showed
substantial iron loss, those younger than 2.0 Ga had practically lost no iron.
It seems that du;:ing this transitional period an abnormally
high amount of organie carbon was buried (Karhu and Holland 1996). This is inferred from a large positive ö 13 C excursion in carbonates deposited during this time interval. If
carbon, accumulated by autotrophie organisms, undergoes
biologically mediated fractionation, carbonates are depleted
of light carbon as common in younger carbonates. Other
results from stable carbon isotopes studies on limestones and
dolomites indicate that substantial amounts of organic carbon
have been buried since ab out 3500 Ma (Schidlowski et al.
1975; Schidlowski 1987). The oxygenreleased bythismechanism was rapidly used up by the processes mentioned
above.
Using the criteria of iron loss from paleosols, the atmospheric oxygen content reached a value of
p02~0.03 atm since sometime between 2.2 and 2.0 Ga
(2200 and 2000 Ma). After this turning point in the
atmospheric oxygen content, the first red beds were
deposited (see below).
Another indicator of an oxygen-poor early atrnosphere are the placer deposits of uraninite (UPs) in
fluvial sandstonesand conglomerates. Such deposits
are older than about 2200 Ma and occur, for example,
in the Witwatersrand in South Africa and in the Elliot
Lake region in Canada (e.g., Pretorius 1981; Holland
1984). In oxidizing conditions, this mineral is unstable
and cannot survlve exposition to weathering and fluvial transport for long.
evolution of life, we have to consider the question
whether and in which way these factors deviated in
pre-Phanerozoic times from the younger Earth's history.
6.5.2 Evolution of the Atmosphere
The Early Atmosphere
Although the atmosphere represents a very small mass
in relation to the total mass of the Earth, it exerts a
dominant influence on the shaping of the landscape,
the existence and evolution of life, weathering of
rocks, transport and deposition of sediments. For this
reason, an atmosphere differing from the present-day
situation must have had a profound impact an all these
aspects.
Ihe evolution ofthe atmosphere was closely related
to the formation of the Earth's crust and hydrosphere.
Most experts hold the opinion that the primordial atmosphere ofthe Earth resembled that of Jupiter. Ihis
planet has sufficient mass to retain light moleeules and
noble gases. In the case of the Earth, however, the
light, volatile components, such as hydrogen and helium, were early lost to space and replaced by an atmosphere mainly caused by outgassing of the Earth's
mantle.
Ihe accretion of planetesimal material, meteoric
impacts, and enhanced radioactive decay raised the
temperature on Earth to such a degree that the primarily more or less homogeneous mass melted. Ihis led to
a differentiation of the accreted mass into core, mantle,
early crust, and a secondary atmosphere derived from
the volatiles of the magma andlor late phases of accretion.
Ihis atmosphere attained a very high density comparable to that of Venus. Its main components were water vapor, carbon dioxide, and nitrogen. Minor constituents included hydrogen, methane, and ammonia.
Light volatiles were continuously lost to space; acids
(HCl, HF) and H 2 S exhaled from volcanoes were
washed out by precipitation. Subsequent rapid heat
loss to space led to cooling of the crust and condensation of water. Evidence for a first hydrosphere on
Earth comes from water-laid sediments ranging back
as far as about 4000 Ma.
The early evolution ofthe Earth, its atmosphere and hydrosphere have been discussed by many workers (e.g.
Schidlowski et al. 1975; Clemmey and Badham 1982; Salop
1983; Holland 1984; Windley 1984; Kempe and Degens
1985; Budyko et al. 1987; Walker andDrever 1988; Wetherill 1990; Condie 1993; Kasting 1993). Further references are
mentioned below.
Chapter 6 Special Depositional Environments
Oxygen in the Early Atmosphere
Whether or not oxygen was present in small quantities
in the early atmosphere is controversial. A limited
amount of oxygen produced by photolysis, i.e., by
dissociation of water vapor as a result of ultra-violet
radiation in the upper atmosphere, was rapidly used up
by oxidation of volcanic gases, ferrous iron dissolved
in sea water (see below), and weathering ofrock-forming minerals. It was not before the plants had started
photosynthesis by using solar energy, water and carbon
dioxide to generate organic compounds and release
molecular oxygen that oxygen could be produced in
larger volumes than were permanently consumed.
However, an excess in free oxygen was only achieved
by the storage of organic marter in sediments. From
then on, the oxygen content of the atmosphere slowly
increased (Fig. 6.18).
Evidence for free oxygen in the atmosphere comes from
paleosols. As long as no oxygen was available, rock surfaces
exposed to the atmosphere lost ferrous iron by leaching.
Later, when the oxygen content ofthe atmosphere had risen,
iron loss became negligible because feme iron cannot be
exported in solution. Of 50 paleosols reported in the literature, 15 could be definitely identified as true palesols (Rhy
and Holland 1998). Paleosols older than 2.2 Ga showed
substantial iron loss, those younger than 2.0 Ga had practically lost no iron.
It seems that du;:ing this transitional period an abnormally
high amount of organie carbon was buried (Karhu and Holland 1996). This is inferred from a large positive ö 13 C excursion in carbonates deposited during this time interval. If
carbon, accumulated by autotrophie organisms, undergoes
biologically mediated fractionation, carbonates are depleted
of light carbon as common in younger carbonates. Other
results from stable carbon isotopes studies on limestones and
dolomites indicate that substantial amounts of organic carbon
have been buried since ab out 3500 Ma (Schidlowski et al.
1975; Schidlowski 1987). The oxygenreleased bythismechanism was rapidly used up by the processes mentioned
above.
Using the criteria of iron loss from paleosols, the atmospheric oxygen content reached a value of
p02~0.03 atm since sometime between 2.2 and 2.0 Ga
(2200 and 2000 Ma). After this turning point in the
atmospheric oxygen content, the first red beds were
deposited (see below).
Another indicator of an oxygen-poor early atrnosphere are the placer deposits of uraninite (UPs) in
fluvial sandstonesand conglomerates. Such deposits
are older than about 2200 Ma and occur, for example,
in the Witwatersrand in South Africa and in the Elliot
Lake region in Canada (e.g., Pretorius 1981; Holland
1984). In oxidizing conditions, this mineral is unstable
and cannot survlve exposition to weathering and fluvial transport for long.