Planet Earth: Time, Origin and Evolution

2.1 Time, Solar System and Early Earth. 1

2.1.1 Introduction. 1

2.1.2 Age of Earth. 1

Early thought 1

The Atom and Radiogenic Dating. 2

2.1.3 Formation of the Solar System.. 5

The Nebula Hypothesis. 5

Chemical composition of the planets. 7

2.1.4 Early Earth. 8

Heating and Differentiation. 8

The Earliest Atmosphere and Oceans. 10

Oxygen in the Atmosphere. 12


2.1 Time, Solar System and Early Earth

2.1.1 Introduction

Earth special place for the development of modern life needs first to be explored in the broader context of planetary formation and Earth’s early evolution, so we start this unit with a brief examination of the origin of our solar system.  Our current knowledge of astronomy has led to an age estimate of our Universe of ~14 billion years.  The evidence is obtained from movements of large objects, such as galaxies, and from evidence of the microwave background radiation.  The red shift of moving objects and today’s temperature of the Universe, among other lines of evidence, led to the theory of the Big Bang.  Within this broad framework, the first question for us is the ages of our planet, which we obtain from radiometric measurements. 

2.1.2 Age of Earth

Early thought

Estimates of the age of earth have occupied us since thought developed.  The topic has scientific and religious components, so we begin by examining early thought.  Recognizing the role of religious thought we refer to 17thC Bishop James Ussher who determined the age of earth through counting generations before Year 0.  Using descriptions of the Old Testament, he calculated that the six days of formation of our world occurred 40xx years before the birth of Christ.  An age of ~6000 years is the literal interpretation of the biblical record, giving one estimate.  Scientists have explored the Earth’s age in other ways, for which we briefly explore early determination and today’s consensus determination.

William Thompson (later Lord Kelvin) used Earth's cooling since its formation from a today’s temperature change with depth to give an estimate of the planet’s age.  Using thermal properties of rocks that were determined in the laboatory and a planetary body that started as a molten mass, he produced and ages in the range of 50-100 my.  This determination was firmly grounded in the physics of late 19th Century, so its results were considered indisputable. Whereas we do not give its derivation, we can readily experiment with Kelvin's calculation. The relationship is: 

                        age = (To - T)2 / ( p*K*GG2),

where To is the formation temperature, T is today's temperature, p is 3.14, K is a material property called thermal diffusivity (we'll use 1mm2/sec) and GG is the Earth's geothermal gradient (25oC/km).  If To is 2000oC, the age of the Earth would be 65 m.y.  Today’s estimate of the early temperature is closer to 7000oC, which would give an age of ~800m.y.  While both estimates are considerable too young, they represent a huge departure for the biblical perspective, leading to clashes among clergy and scientists.  At the time it was hard to argue with sound physics, until a major new discovery was made around the turn of the century: radioactivity.  This property led to today’s determination for the Earth’s age of 4.56Ga.

The Atom and Radiogenic Dating

Atom = nucleus (+ charge) and electrons (- charge)

Nucleus = protons (+ charge) and neutrons (0 charge)

Protons, neutrons = quarks bound by gluons

# protons: atomic number (element, e.g., 2He)

# neutrons+# protons: mass number (e.g., 4He)

Isotopes = element with different masses; 
            variable # neutrons

rnucleus is 10E-11cm  ratom is 10E-7cm

Atom is10,000 larger than nucleus! (1cm vs. 100m)

Mass proton/neutron is ~10E-27kg; mass electron is ~10E-30kg

So, atom mostly empty space and most mass in nucleus.


Up to about the element silica (Si) in the periodic Table of the Elements, the number of protons in an element equals the number of neutrons.  Heavier elements can have several isotopic numbers, meaning different numbers of neutrons, but the same number of protons.  For example, the element Rubidium has the isotopes 85/37 Rb and 87/37 Rb.  The discovery of radioactivity was that the occurrence of some isotopes is unstable, such that a new element is formed spontaneously. Of the two Rb isotopes, 87/37 Rb is unstable and it changes to the element Strontium (87/38 Sr) by the conversion of a neutron into a proton and an electron. The electron is expelled from the nucleus of the new element, which produces a dangerous and now familiar side effect, called radiation.  This particular type of radioactive decay is called beta decay (b), but there are several other types of radioactive decay, which are illustrated in the Figure X. 


Figure X.  Types of radioactive decay.

Alpha decay is the emission of particles that contain two protons and two neutrons (He).  This results in a daughter with a lower atomic number (-2) and a lower mass number (-4).  Beta decay describes the emission of an electron, which converts a neutron into a proton.  The atomic number increases by 1, whereas the mass number remains the same.  Another form of beta decay is when a nucleus catches an electron, resulting in the conversion of a proton to a neutron.  This electron capture process, results in a decrease in atomic number, but no change in mass number.  Gamma decay (g) produces gamma rays, which is electromagnetic radiation from photon emission.


            In the language of radioactivity, rubidium would be called the parent isotope and strontium the daughter isotope. The number of isotopes that decay per unit time is proportional to the total number of parent isotopes present. A convenient measure to express this property is through the concept of the half-life (t˝) of an isotope. The half-life is the time required for half of a given number of parent isotopes to decay to a daughter isotope.  Plotting this behavior shows the non-linearity of the relationship.

Figure X.  The half-lives of an element.


            We used the Rb-Sr system as an example, but many radioactive isotopes exist.  Table X lists common radiogenic systems with their corresponding half-lives and decay constants.  Note, for example, that it takes nearly 49 billion years to change 50% of Rb atoms into Sr atoms.


Table X. Commonly Used Long-Lived Isotopes

Parent (P)

Daughter (D)

Reference (S)

Half-life, t˝
(109 y)

Decay constant, l






































Note: * 204Pb is not stable, but has an extremely long half life of ca. 1017 years.


A useful analogy to illustrate the fundamentals of geochronology is an hourglass.  If we start with one side of the hourglass full (containing the 'parent') and the other side empty (collecting the 'daughter'), we only need to know the rate at which the sand moves from one chamber to the other (represented by the half-life and the decay constant) and the amount of sand in the daughter chamber or the amount of parent remaining to determine how much time has passed.  However, in reality matters are slightly more complicated.  At the time the radiogenic clock starts ticking, most natural samples already contains some daughter material; in other words, some sand is already present in the daughter chamber. This amount of daughter is referred to as the initial daughter. So, when we measure the amount of daughter product in our specimen we are in reality combining the amounts of daughter from decay of the parent, called the radiogenic daughter, and initial daughter. We resolve this obstacle by separately analyzing constituents of the material that contain different amounts of 87Rb.  By plotting these results we can extract the formation age of the material, without interference from initial daughter.

Applying the method to objects from the earliest times of our solar system, chondritic meteorites, we obtain a remarkably consistent result of 4.56 billion years for the age of these objects and by inference of larger planetary bodies (Figure X).  Radiogenic age measurements on rock and minerals from Earth are not that old. The oldest rock, found in northern Canada, is about 4 Ga, whereas the oldest mineral is found in NW Australia and is about 4.3 Ga.  Samples collected through the lunar program of the late 1960s and early seventies gave older ages. The first moon rock picked up was later dated at 3.6 billion years!  All moon rocks examined to date are in the range 3.1 - 4.6 billion years. 

Figure X. Radiometric ages of various meteorites that converge on 4.56Ga.


            Comparing the age of the Universe and our Solar System leaves the important realization that our Sun and planets were not formed at the time of the Big bang.  Rather they were formed from recycled matter.  We first turn to the reigning hypothesis of solar system formation before returning to the earliest history of our planet.


2.1.3 Formation of the Solar System

The Nebula Hypothesis

The formation of planets rotating around Sun is described by the Nebula Hypothesis, which involves rotation and turbulence.  The probable sequence of steps in the formation of the solar system is illustrated in Figure X. 


Figure X. The Nebula hypothesis. (a) Gravitational contraction of a rotating gas cloud leads to a dense central region (eventually forming the Sun) and a more diffuse, flattened nebula. (b) Dust particles from the nebula settle onto a disc. (c) Accretion of dust into numerous small planetesimals. (d) Eventually larger bodies capture the smaller ones.


Gravitational contraction of a slowly rotating gas cloud leads to a denser central region, eventually forming the Sun, and a more diffuse, flattened outer region.  Increasingly dust particles from the nebula settle onto the flattening cloud, which takes the form of a disc.  Accretion of dust led to the formation of numerous small planetary objects, or planetesimals, each a few kilometers in diameter.  The inevitable collisions between these planetesimals lead to capture or disintegration, and more deflection of their orbits.  The continued collisions eventually creating larger bodies, called protoplanets, through capture of smaller ones.  Uncondensed gas in this cloud is meanwhile blown away by the "solar wind", leading to today’s subdivision among planets into inner, also called rocky or terrestiral planets and outer, or gaseous (also called jovian) planets (Figure X), which characteristic geochemistry.


Figure X. Relative sizes, sequence and subdivision of the planets in our solar system 


Chemical composition of the planets

The outer planets (Jupiter, Saturn, Uranus, Neptune) have compositions different from Earth and its neighbors and more consistent with the composition of the solar system: lots of hydrogen and helium. Table X compares some properties of the inner and outer planets.


Table X. Some Properties of the Planets


Diameter (km)

Distance from Sun 
(x106 km)

Surface temperature


Main atmospheric constituents











5.4 (rocky)






5.3 (rocky)






5.5 (rocky)

N2, O2





3.9 (rocky)






1.3 (icy)

H2, He





0.7 (icy)

H2, He





1.3 (icy)

H2, CH4





1.7 (icy)

H2, CH4








The reason for the difference between the rocky and dense inner planets, and the icy/gaseous outer planets is determined by the type of material that forms, or condenses, in the solid form given the temperature of the particular part of the Nebula.  At temperatures above about 1300K, metals and silicates can condense and become solid dust grains, while at lower temperatures more volatile minerals become solids.  At temperatures of less than ~400K, hydrogen-bearing gases such as methane and ammonium become solids, but hydrogen and helium remain gases.  For the inner planets, at high temperatures because they are close to the Sun, the planet-building dust grains were therefore made up of materials such as silicates and iron, forming the rocky planets.  This scenario is called condensation theory and is illustrated in Figure X.


Figure X.  Condensation sequence of minerals in the solar nebula as a function of temperature. At high temperatures only metals and silicates can condense and become solid dust grains. At lower temperatures more volatile phases become solids.


What Happened to the Earth's H and He?  Initially Earth formed with rocky grain accretion and with an abundance of H and He.  The Earth's gravity was so small that the lightest gases escaped out of the atmosphere to space. The same process occurs today with helium and hydrogen.  The Earth was part of Nebula that was hot enough for hydrogen-bearing gases to remain as gases (and not condense). The solar wind simply swept these gases away. The hydrogen and helium would have been blown away by the solar wind.  For the outer planets, farther from the Sun, the hydrogen and helium was retained by a combination of the large gravity for these enormous bodies and the formation of ice. 


2.1.4 Early Earth

Heating and Differentiation

Based on observations from seismology, some of which we’ll revisit later in thus unit, we learn that Earth's solid body is composed of several layers of varying composition and density.  The Earth's core is composed of two layers, an inner core of solid iron and an outer core of molten iron.  Above the core lies the mantle, which is made up of dense silicates, and the crust, which is the outermost, thin  layer of the solid Earth.  Differentiation in the first few 100's of millions of years led to the formation of the core and the mantle, and allowed the escape of gases from the moving interior, eventually leading to the formation of an early atmosphere and oceans. 

The earliest Earth was an unsorted conglomeration, mostly of silicon compounds, iron and magnesium oxides and smaller amounts of other natural elements.  The protoplanet became increasingly hotter as it grew by the actions four different effects: 

  1. Accretion. Impacting bodies bombard Earth and convert their energy of motion (kinetic energy) into heat.  The collision with one very large object was responsible for the "extraction" of the Moon from Earth. 
  2. Differentiation.  The gravitational potential energy from dense, sinking objects converts into heat during core formation.
  3. Self-compression. As Earth gets bigger, the increasing gravity forces the mass to contract into a smaller volume, producing heat.  This is analogous to a bicycle pump that gets hotter on compression. 
  4. Short-lived radiogenic isotopes.  The protoplanet absorbed the energy released in radioactivity from elements with short half-lives, such as XX and XX, heating the planet.  Today’s radioactivity is a much slower, but steady source of heat.

At some point within the first few hundred million years of Earth, the outer part down to a depth of about 500 km became so hot that iron started to melt.  The molten iron collected and began to sink under its own weight.  About one third of the primitive planet's material sank toward the center, and in the process, heating rates increased and most of the planet became liquefied with a temperature around 7000K.  Rather than water, Earth’s earliest surface ocean was filled with magma.  The formation of a molten iron core was the first stage of differentiation of Earth, at which it was converted from a homogenous body, with roughly the same kind of material at all depths, to a layered body, with a dense iron core and a thick silicate mantle.  Subsequent differentiation separated the a thin crust, composed of light silicates with relatively low melting points, from the denser silicate mantle.

Let’s look at this scenario from an elemental perspective.  Figure X shows the elemental abundances of Earth.  The main elements, Fe, Si, Mg and O react to form relatively light silicate compounds, comprising SiO4, which consumes most of the O atoms.  The remainder O reacts with Fe and S to form oxides, after which Fe.  These heavier Fe atoms form Fe droplets that sink to the center of the planet, producing the core.  The silicates and oxides form the mantle layer.  Other elements such as aluminum, calcium, potassium, and sodium are far more abundant in the crust than in the whole Earth, because they form lightweight chemical compounds, which are easily melted.  Lava flowing from the partially molten interior spread over the surface and solidified to form a thin crust. This crust melted and solidified repeatedly, with the lighter compounds moving to the surface.  The end result of these processes was a continental land mass that would grow over time, in first two billion years of Earth. 

Figure X.  Relative proportions of Earth’s common elements, by weight and by atoms.

As a result, the earth evolved from a homogeneous body into a layered body, with the main properties in Table X.


Average Thickness (km)

Average Density (x103 kg/m3)

Fraction of Total (%)

Principal Constituents





N2, O2










Silicates and other oxides





Mg silicates






Fe, +/-S (liquid)

Fe-Ni (solid)








The Earliest Atmosphere and Oceans

The convective process associated with a differentiating planet is responsible for yet another key component of our planet: an atmosphere.  As planetary material violently overturned, volatile gases, such as CO2 and H2O-vapor are released, accumulating in a gaseous surface layer.  This outgassing produced our earliest atmosphere, in which reactions produced other basic compounds such as methane (CH4) and ammonia (NH3).  While far less abundant, the gases methane and ammonia are highlighted as they are the key compounds of amino acids, which are the fundamental building blocks of life’s proteins.  This has been illustrated by the classic Miller-Oparin experiment, where these gases were combined with electrical charges, representing lightning, to show the formation of more complicated compounds.  These experiments did not form long protein-like molecules, so the origin of life remains debated.  Today, billion of years later, the outgassing process continues as volcanism, albeit much less dramatic, releasing mainly H2O, CO2, SO2 and lesser amounts of gases such as N2.

Figure X. Volcanic outgassing.


With H2O the dominant species in the outgassing process, earth’s early atmosphere became saturated with water vapor, leading to the beginning of an era of continuous rain on the planet.  This change marks conditions on our planet that are distinct from our neighbors and critical for its subsequent evolution.  As the planetary surface cooled and vapor saturation is reached in an early atmosphere, H2O can accumulate as liquid or ice, or remain as a vapor, depending on the surface temperature. 

Phases of H2O as a function of atmospheric vapor pressure and surface temperature.


A phase diagram of water (Figure X) shows these scenarios for Mars, Earth and Venus.  As vapor increases in a Martian atmosphere, its relatively low temperatures will result in the formation of ice.  Earth’s surface temperature, on the other hand, allows the formation of H2O in liquid form and ice only at the colder poles.  Vapor saturation of H2O on Venus is never reached, so the atmosphere continues to accumulate H2O.  On Earth, torrential rains lasting millions of years gradually accumulated in increasingly larger ponds that further cooled the planet and eventually led to the first oceans.  Once liquid water became available at the surface, the atmospheric composition further changed, as other atmospheric species such as CO2 are soluble in water.  Lowering the CO2 content of the early atmosphere meant lowering the natural greenhouse effect, further lowering the Earth’s surface temperature toward conditions suitable for life to evolve.  This transition marks another major event in our early history, as it ultimately leads to the presence of atmospheric O, which is key to modern life.

Oxygen in the Atmosphere

Today’s atmosphere contains mostly N2 and 21% oxygen, yet no free oxygen was produced in the early years of planetary differentiation and outgassing in contrast to N2.  The origin of initial O is linked to breakdown of water vapor from sunlight and particularly the formation of early life, which is one example of the natural feedback between the biotic and abiotic world.  As the atmosphere accumulates H2O and CO2, reaction with solar radiation produces oxygen in the upper atmosphere; 2H2O to 2H2 and O2; a process called photolosys.  This oxygen further reacts to form a special oxygen compound called ozone, O3, which produces a radiation shield for the sun’s UV light.  Once the shield was in place, life became shielded from these high-energy radiation, and became a productive source of O, eventually producing today’s high O content

The main energy capture and release mechanisms of life consist of reduction and oxidation reactions, also called redox systems.  Oxidation is removal of electrons, which releases energy.  A warm campfire, but many oxidation reactions occur.  In the case of the element Fe, oxidation (or rusting) change its charge from Fe2+ to Fe3+.  Reduction is the addition of electrons, which consumes energy.  Analogously it changes the Fe charge from Fe3+ to Fe2+.  We focus on iron as its charge is key to Earth ancient record of an oxidizing world.  But, familiar redox reactions in the biotic world are Photosynthesis, a reduction reaction:

CO2 + H2O + e- → CH2O + O2; CO2 is the electron acceptor; and

Respiration, an oxidation reaction:

CH2O + O2 → CO2 + H2O + e-; CH2O (formaldehyde) is oxidized by oxygen

Many reducing substances (or electron donors) dissolve in water, so their preservation in ancient rocks are indicative of this condition.  The presence of reduced Fe in banded-iron formation indicates an initially O-free atmosphere.  These rocks, which are found in many locales older than 2.5 billion years, contain the mineral magnetite, Fe3O4, which is the reduced form of Fe-oxide.  The appreance of redbeds around 2billion year, containing hematite (Fe2O3), the oxidized form of Fe, signals that free O has entered the atmosphere.  The main source for atmospheric oxygen is photosynthetic aquatic organisms called “blue-green algae” which are found in rocks as old as 3.8 Ga.  These organisms are not eukaryotes (meaning their cells have a nucleus), like today’s algae, but are relatives of bacteria (prokaryotes, meaning cells without a nucleus) and chloroplast in plant/algae cells.  For a long time, the oxygen produced did not build up in the atmosphere, as it was taken up Banded Iron Formations (BIFs; picture) and continental red beds.  It was not until about 1.5 billion years ago that the reservoirs of oxidizable rock became saturated and the free oxygen stayed in the air.  Once oxygen had been produced, the sun’s ultraviolet light split the molecules, producing the protective ozone UV shield as a by-product.  Only at this point could life move out of the oceans and did respiration evolve.



Figure X.  Banded-iron formation, fossilized stromatolite and today’s blue-green algae in Australia’s Shark Bay.


            Today, blue-green “algae” are rare, but one example can be found in Australia’s Shark Bay (Figure).  The rock record has many location where fossilized remnant are preserved, called stromatolite (Figure ).  Life started to have a major impact on the environment once these photosynthetic organisms evolved.  They fed off atmospheric carbon dioxide and converted much of it into marine sediments consisting of the shells of sea creatures.

Figure X.  Cumulative history of O2 by photosynthesis over geologic time.