April 29, 2011
The Origin of the Earth
Part I: Introduction, the Scientific Toolbox, and Cosmic Starstuff
The Cosmos is all that is or ever was or ever will be. Our feeblest contemplations of the Cosmos stir us; there is a tingling in the spine, a catch in the voice, a faint sensation, as if a distant memory, of falling from a height. We know we are approaching the greatest of mysteries; in the last few millennia we have made the most astonishing and unexpected discoveries about the Cosmos and our place within it; they remind us that humans have evolved to wonder, that understanding is a joy, that knowledge is prerequisite to survival. I believe our future depends on how well we know this Cosmos in which we float like a mote of dust in the morning sky.
-Carl Sagan, Cosmos
The study of the origin and development of the early Earth is one of the most intriguing and important research topics in science today. Understanding Earth’s history in the context of the larger Cosmos is both awe-inspiring and humbling. Earth is but a tiny speck in the solar system and is dwarfed by the greater Cosmos. Yet, Earth is our home and harbors the only life of which we know. So, to us, study of the Earth– both past and present-– is fascinating and also essential for our survival as a species. Psychologically, there is also something especially powerful about studying the early Earth. Just as we value learning about our ancestors and family histories, we value learning about the history of our planet. More practically, a better understanding of early Earth may provide information on how our planet will respond to stresses such as pollution, greenhouse gas emissions, and magnetic field reversal. Furthermore, if humans ever colonize planets in other solar systems, knowing what conditions and processes lead to an Earth-like planet may help scientists locate habitable planets. Study of early Earth is also relevant because today, at least in America, religious creationism has a strong hold in many places. Better understanding Earth’s long history will aid those who confront the creationists and fight to teach valid science in our school systems.
The goal of this paper is to summarize the processes that scientists believe, to the best of their current understanding, led to the formation and development of the Earth. Because Earth formed from stardust, this paper will begin with an overview of nucleosynthesis, which occurs primarily in stars. The processes leading to the formation of various elements in the Cosmos will be described and trends in elemental abundances explained. Much information about Earth’s history can be gleaned from a comparison of the abundances of elements in the Cosmos and solar system with the abundances of elements on Earth. Therefore, differences in the composition of the Cosmos, the sun, and the Earth will be noted and comments made about how elements are fractionated and concentrated in the solar system. Similarly, a general discussion of the solar nebula and the early development of the sun and solar system will be provided in order to put the Earth in a larger context.
Next, a brief overview of meteorites will be provided as Earth is believed to have accreted from planetesimals with compositions similar to meteorites, in particular carbonaceous chondrite meteorites. Understanding why and how geologists use meteorites to estimate the intial composition of the Earth is crucial to understanding the assumptions inherent in any geophysical or geochemical model using chondrite values for the composition of the early Earth. Following this discussion of meteorites, the accretion of the Earth from planetesimals will be discussed. The importance of late-stage collisions between large planetesimals will be emphasized. At the same time, the origin of the Moon from Earth via a Mars-sized impactor will be examined. Theories about a magma ocean that may have formed as a result of this giant impact will be presented as this magma ocean may have played a significant role in the differentiation of Earth’s iron core.
Finally, the formation of Earth’s core and the subsequent differentiation of the upper Earth into an enriched continental crust and a depleted mantle (which now melts and forms oceanic crust) will be described. Possible evidence for an early protocrust, supposedly subducted deep in Earth’s mantle, will also be analyzed. Clearly, discussing all these topics related to the origin and development of the Earth is a large task for a small paper. However, this paper will hopefully at least provide a sense of how scientists are able to decipher the early history of our home planet and a general overview of Earth’s origin and history.
How Scientists Study Early Earth:
In order to appreciate the story scientists have reconstructed concerning the origin and development of Earth, one must first have an understanding of how scientists are able to glean information about the early Cosmos. What are the objects that scientists study in order to learn about the early Earth, solar system, and universe? Scientists start with objects close-to-home: the Earth and her solo moon. The Earth is a terrestrial, differentiated planet with active plate tectonics and an atmosphere, hydrosphere, and–uniquely so far as we know– a biosphere. The moon is smaller and more barren but is closely related to the Earth. Next, scientists venture further and study the sun and other planets in our solar system. Scientists can also study asteroid belts and meteorites, which are small pieces of asteroids and planets that fall to Earth. Scientists can also glean information from objects outside the solar system. Scientists study galactic cosmic rays and also remotely observe other stars, solar systems, and solar nebulae.
In addition to appreciating the objects scientists study to learn about the origin and development of the Cosmos, one must also appreciate the tools which scientists use to study these objects and to develop theories about the origin of the Cosmos, our solar system, and the Earth. There are three important categories of tools scientists use to study the early Cosmos: spectroscopy, cosmochemistry, and computer modeling. Spectroscopy is the study of the interaction of electromagnetic radiation and matter. Study of absorption lines created by the interaction of radiation and matter can be used to determine the chemical composition of stars, solar nebulae, and cosmic rays. The cosmochemical toolbox includes major elements, trace elements, and isotopes. Within isotopic analysis there are short-lived isotope systems, which are particularly useful for constraining the time of various events in the solar system, and long-lived isotope systems. Computer models are also widely used to simulate conditions in the early solar system and Earth. Computer models are very useful for visualizing processes in the early Cosmos and determining their likelihood, but one must carefully understand the assumptions of the models and must work towards constraining the parameters of the models.
Cosmic Abundances of the Elements and Nucleosynthesis:
From the study of stars, galactic cosmic rays, and chondrite meteorites, scientists have determined the cosmic abundances of the elements (Figure 1).
Looking at a plot of abundance verses atomic number (Z), one can make several observations about the chemistry of the Cosmos:
1. Hydrogen (H) is by far the most abundant element, ~75% of the Cosmos by weight (Dickin, 2006).
2. Hydrogen (H) and helium (He) together comprise ~99% of the Cosmos by weight (Dickin, 2006).
3. Generally, heavier elements are less abundant than lighter elements.
4. Certain elements are depleted relative to their neighbors. In particular, the light elements lithium (Li), beryllium (Be), and boron (B) are especially depleted.
5. Certain elements are enriched relative to their neighbors. Most notable among these are iron (Fe), particularly the isotope 56Fe (Dickin, 2006), and lead (Pb).
6. The elements in general have a sawtooth pattern with elements alternately enriched and depleted with atomic number.
As we will see, these observations can be explained by looking at how elements, and stable isotopes in particular, are made. One must understand the various processes and stages of nucleosynthesis as well as understand which elements are stable and store the most potential energy. Perhaps the easiest of the above observations to explain is the characteristic sawtooth pattern of the elements. Generally, isotopes with an even number of protons and an even number of electrons are energetically favored over isotopes with odd numbers of protons and neutrons. By far, most stable isotopes have an even number of protons and neutrons and an overall even atomic number (protons + neutrons). Thus, the alternating even-odd pattern of abundances can be explained by the fact that even atomic numbers are favored and thus more abundant than odd atomic numbers.
The next major observation that can be explained is the abundance of H and He. H has one proton (and sometimes one neutron and one electron) and is the easiest element to form because of this. He is essentially just two hydrogens put together and has two protons, two neutrons, and two electrons generally. Together, these two elements were created in large quantities in the big bang, an explosive event ~13.7 billion years ago in which the universe was created from an initially extremely hot, dense state (Norton, 1994). During the big bang, there was rapid expansion during which free quarks and gluons condensed into larger particles and the first elements: H, He, and a small amount of Li (Dickin, 2006).
Beyond H, He, and Li– the three lightest elements– larger elements must be produced by synthesis in stars and through a process called spallation. The elements from carbon to calcium are produced by nuclear fusion in a star. Elements heavier than carbon are produced by neutron capture (either the s-process or r-process), proton capture, or e-process interactions between nuclei and free protons and neutrons (Dickin, 2006). S-process or slow-process neutron capture produces about half of the elements heavier than iron and occurs at moderate neutron density and temperature conditions in middle-aged stars. The e-process occurs at high-temperatures in stars just before they explode as supernovae. The e-process produces elements such as iron and other first-series transition elements (Dickin, 2006). R-process or rapid-process neutron capture and proton capture occur in supernovae and produce unstable isotopes which rapidly decay back to more stable isotopes.
Certain light elements can also be produced through cosmic ray spallation, a process that produces elements through the fission of heavier elements. Spallation occurs when an element is bombarded with a cosmic ray (Dickin, 2006). Cosmic rays are primarily a high-energy proton stream, and when a heavy nucleus is hit with a high-energy proton it may release nucleons and form lighter elements. Spallation is the primary way in which the light elements Li, Be, and B are produced. Figure 2 relates atomic number with various nucleosynthetic processes. Note that Figure 2 is a simplification as many elements are produced by more than one process. Figure 3 shows the general evolution of a star and the times at which various nucleosynthetic processes occur.
Figure 3: The Evolution of a Star and Nucleosynthetic Production. Figure from Dickin, 2006.
Returning now to an explanation of the cosmic abundances of the elements, the general abundance of lighter elements relative to heavier elements can be explained by the fact that heavier elements take longer and are harder to make in stars as you must keep adding neutrons to a nucleus. The depletion of Li, B, and Be can be explained by the fact that these elements are bypassed in stellar fusion because the nuclear binding energies of these elements are very low. Fe, on the other hand, is at the peak of the nuclear binding energy curve, which is shown in Figure 4. Nuclear binding energy is the energy required to dissemble a nucleus into individual protons and neutrons and can be thought of as the energy a nucleus can store. Nuclear systems that are bound have less potential energy than unbound systems and are favored. Thus, elements with high binding energies (such as iron) are favored over elements with lower binding energies because such elements can store more potential energy.
One last comment about the cosmic abundance of the elements regards neutron “magic numbers,” which lead to abundances of isotopes with certain numbers of neutrons such as 50, 82, and 126 (Dickin, 2006). The abundance of these isotopes has to do with something called neutron-capture cross-section, which is a measure of how readily a nucleus can absorb a neutron (Dickin, 2006). Elements with low neutron-capture cross-sections will not be converted to species of higher atomic mass as easily and so will be more abundant relative to their neighbors.
Part II: Crustal Chemistry, the Solar Nebula, and the Solar System
Terrestrial Abundances of the Elements:
After one has a basic understanding of the cosmic abundances of the elements, one may notice that the terrestrial planets, Earth included, have a chemical composition that differs significantly from the cosmic chemical composition. The differences between the cosmic and terrestrial abundances of the elements can provide important clues about the processes of Earth’s formation and evolution. Earth’s crust is mostly oxygen (about 47% by weight) and also contains significant amounts of silicon, aluminum, iron, calcium, sodium, potassium, and magnesium as shown in Table 1 (Lutgens and Tarbuck, 2003).
The reasons why certain elements are depleted or enriched in Earth’s crust will be explained further in subsequent sections, but for now it is worth bearing in mind that elements can be either volatile or refractory. Volatile elements are those elements with low boiling points and are easily vaporized. Refractory elements, on the other hand, have high boiling points and are fairly robust when heated. Earth’s elements can also be thought of as generally being atmophile, lithophile, siderophile, or chalcophile (see Figure 5). Atmophile elements (e.g. H, He) are highly volatile and are concentrated in Earth’s atmosphere and easily lost to space. Because of their volatility, atmophile elements are depleted in the Earth relative to their cosmic abundances. Lithophile elements have a strong affinity for oxygen and are concentrated in the silicate portion of the Earth (i.e. the crust and mantle). Siderophile elements, which have an affinity for iron, and chalcophile elements, which have an affinity for sulfur, are concentrated in Earth’s iron core.
The Solar Nebula:
The solar nebula was a dense, rotating cloud of interstellar gas and dust that collapsed to form our solar system. Scientists gain information about what our solar nebula may have looked like from the composition of the sun, our star, and also by studying other nebulae. In the Milky Way Galaxy there are many nebulae scientists can observe. In general, nebulae are dozens of light years wide and are composed of gases, mostly hydrogen but also some nitrogen and oxygen as well as tiny dust particles, which are mostly carbon, silicates, and iron that were produced in stars (Norton, 1994). These dust particles may also be encased in layers of methane, ammonia, and water ice (Norton, 1994). Significantly for the origin of life, some of the densest nebulae are observed to contain organic compounds of varying complexities (Norton, 1994).
Our own nebula likely consisted of approximately 98% hydrogen with 2% heavier elements (Dickin, 2006). This composition indicates that the solar nebula formed from the ashes, so to speak, of dead stars as these heavier elements had to be produced through stellar processes. As the well-known astronomer Carl Sagan was fond of pointing out, we are all made of stardust as is everything on our planet and in our solar system. In general, nebulae with higher abundances of heavier elements produced in stars are more likely to contain terrestrial planets such as Earth (Dickin, 2006).
The Development of the Solar System:
The solar system began forming when the solar nebula started collapsing gravitationally. What triggered the collapse of the solar nebula is debatable, but one likely cause is a shockwave produced by a supernova explosion during the death of a nearby star (Norton, 1994). As the solar nebula collapsed gravitationally, three important developments occurred: the solar nebula became hotter, it began to spin quickly, and it flattened into a disk (Norton, 1994)
The solar nebula became heated as gravitational potential energy was converted into kinetic energy. Also, the solar nebula became heated because as particles fell into the gravitational well, they moved more quickly as the gravitational forces became stronger because the particles were closer together.
The solar nebula began spinning quickly because of the conservation of angular moment. When the solar nebula was a spherical gas cloud before collapse, the nebula was rotating very slowly and had a direction of net angular momentum. Angular momentum is defined as L = Iw where w is the angular velocity of the object, and I = mr^2 where m is the mass of the object and r is the perpendicular distance from the object to the axis of rotation. As objects move closer to the axis of rotation about which they are rotating, they move faster. Think of a spinning ice skater for a simple example: as the skater pulls his arms closer to his body, he spins more quickly. Similarly, as the solar nebula collapsed, the particles had to move faster in order to conserve angular momentum. Finally, the solar nebula flattened into a disk because the direction of net angular momentum was favored and so the particles began preferentially moving in the same direction.
Eventually, the central, dense center of the collapsing solar nebula reached the pressure and temperature conditions under which nuclear fusion could begin and our star the sun was born. Currently, our sun is a middle-aged star main sequence star steadily burning hydrogen and releasing energy through nuclear fusion. When the sun was first born, however, it was a T-Tauri star (see Figure 6), a young star which is larger, less hot, and had a much more powerful solar wind than a main sequence star (Norton, 1994). One important point to note is that the powerful solar wind associated with this T Tauri star would have blown away other gases in the developing solar system. Thus, any planets or planetesimals forming from these gases would end their growth once the powerful solar wind of the T Tauri protosun developed.
As the sun was forming, the protoplanetary disk around the sun was cooling and condensing. Dust (metals, silicates) and ice (water, methane, and ammonia) condensed out of the solar nebula and began sticking together to form larger particles (Norton, 1994). Some of the particles gradually began growing larger by collisions which led to the capture of other particles. Eventually, the largest particles began growing very large in a process known as oligarchic or runaway growth (Wetherill and Stewart, 1993). The larger the objects became, the faster they would grow because these large objects were moving more slowly and had larger volumes, making collisions with other particles more likely. Additionally, very large growing objects could also gravitationally capture smaller, nearby particles. From condensation and oligarchic growth, fairly large (up to Mars-sized, perhaps) planetesimals could form (Canup and Agnor, 2000). However, as will be described below the production of a large terrestrial planet such as Earth seems to require giant collisions between large planetesimals.
There are three main types of planets in our solar system: terrestrial, gas giant, and icy (see Figure 7). The terrestrial planets are Mercury, Venus, Earth and Mars. The gas giant planets are Jupiter, Saturn, Uranus, and Neptune. Further out, there are some icy planetesimals such as Pluto. Because most of the solar system is hydrogen and helium, terrestrial planets with concentrations of heavier elements can only form close to the sun where the temperatures are too warm for hydrogen compounds to condense out of the solar nebula, so only heavier, less volatile elements condense. Farther from the sun, where temperatures are cooler then about 150 K, hydrogen compounds are able to condense and gas giants are able to form. Because there is so much more material available for these gas planets to form since hydrogen is so abundant, the gas planets become giants, capturing gases from the solar nebula because they are so large. One the T Tauri solar wind sweeps away the solar nebula gases, however, the growth of the gas giant planets is stopped. Finally, at the very furthest, coldest reaches of the solar system icy planets, such as Pluto, form.
Part III: Rocks from Space and the Accretion of the Earth
A meteorite is a rock from space that survives passage through Earth’s atmosphere and lands on Earth’s surface. Objects that become meteorites on Earth’s surface are called meteoroids while still in space and are referred to as meteors when falling through Earth’s atmosphere (Norton, 1994). Meteorites originate from various places in the solar system. Primarily, meteorites are fragments of asteroids from the asteroid belt between Mars and Jupiter. Potentially, some meteorites are also fragments from farther asteroid belts or from asteroid belts that no longer exist in the solar system. Some meteorites have also been identified to originate from the moon and other planets in our solar system. Currently, there are 34 meteorites identified as originating from Mars (NASA’s Mars Meteorites webpage). Most meteorites from Mars probably were chipped off the planet’s surface during a meteorite impact and flung out of the planet’s gravitational influence into space heading towards Earth (NASA’s Mars Meteorites webpage). However, the parent bodies of most meteorites are unknown and may no longer even exist, having been broken into fragments. Regardless, families of meteorites can be identified based on chemical composition. Oxygen isotopes are especially useful in defining families of meteorites (Norton, 1994). Several thousand meteorites have been discovered on Earth, and many more are discovered each year, mostly in deserts or in Antarctica, places where there are few terrestrial rocks and where there are active searches for meteorites.
There are three general categories of meteorites: iron meteorites, stony meteorites, and stony-iron meteorites. Iron meteorites are primarily iron and nickel and are believed to represent the iron cores of differentiated asteroids (Norton, 1994). Since iron meteorites come from the cores of asteroids, the parents of this group of meteorites must have been broken up thoroughly and likely no longer exist. There are two categories of stony meteorites: chondrites and achondrites. Chondrites, which will be described in detail below, are believed to be the most primitive material in the solar system (Carlson and Lagmair, 2000). That is, chondrite meteorites represent undifferentiated asteroids. Achondrite meteorites, on the other hand, are believed to represent igneous material, such as basalt, originating on a parent body with a core and differentiated upper section. Lastly, stony-iron meteorites are believed to come from the core-mantle boundary of a differentiated asteroid. These unusual meteorites are composed of crystals, often olivine, in an iron-nickel matrix. Figure 8 shows where various meteorites are thought to originate in an asteroid and compares these locations with a larger, differentiated planet.
The most important meteorites to study for the origin of the Earth are chondrite meteorites. These meteorites are believed to represent very primitive asteroids that have undergone very little differentiation, heating, or metamorphism since they formed at the beginning of the solar system. More than 85% of known meteorites are chondrites, which suggests that most of the solar system’s rocky matter is undifferentiated and that differentiated bodies such as the Earth are relatively rare (Norton, 1994). Chondrites often contain chondrules, which are small, spherical, silicate inclusions. There is some debate about how chondrules form, but most likely they formed from the initial heating of early solar system dust and thus represent very old material (Norton, 1994; Carlson and Lugmair, 2000). Many chondrites also contain inclusions called CAIs which stands for calcium-aluminum inclusions. CAIs are large, irregular, white inclusions that likely represent the first material that condensed from the solar nebula (Carlson and Lugmair, 2000).
Chondrites are generally classified based on their texture and chemical composition (see Figure 9), both of which are related to the degree of alteration and thermal metamorphism of the meteorite. The least altered and metamorphosed chondrites are carbonaceous chondrites. Carbonaceous chondrites are rich in organic compounds and water-bearing minerals, which can only exist in a meteorite that has not been heated significantly (Norton, 1994; Carlson and Lugmair, 2000). The most primitive of the carbonaceous chondrites are the CI chondrites of which there are only five samples (Norton, 1994). CI chondrites are believed to have not been heated about 50º C and contain significant water (up to 20%) and organic molecules such as hydrocarbons and amino acids (Norton, 1994). These CI chondrites likely grew in the cooler, outer solar nebula and represent the most primitive solar system material ever found. Thus, chondrites and CI chondrites in particular are believed to represent the original, undifferentiated composition of the Earth.
An initial chondritic composition for the Earth is assumed in many geochemical and geophysical models of the formation and differentiation of the Earth. That Earth accreted from chondritic planetesimals is thus an important assumption made in all of these models. However, just how valid is this assumption of the chondritic Earth? There are several potential problems with this assumption. First of all, chondrites– even carbonaceous chondrites– have varying compositions. The assumed chondritic compositon of the early Earth is actually the average composition of many individual chondrite meteorites. There is some question as to just how representative this average composition may be since it is based upon chemical analysis of the few chondrites which have fallen to Earth; there are only five known CI chondrites, for instance. If there were to be another big meteorite fall, there is a possibility that the “average” composition of chondrites could change.
Furthermore, the oxygen isotopic composition of chondrites indicates that there is considerable variation in the parent bodies of these meteorites (Carlson and Lugmair, 2000). Thus, averaging chondritic compositions may not be fair as Earth may have accreted from a specific family of chondrite planetesimals. That chondrites have varying compositions is no surprise as, very likely, the solar nebula was heterogeneous. Indeed, spectrographic observations of other nebula confirm that nebulae are, in general, heterogeneous.
Another concern is the difference in ages of CAI and chondrule inclusions in chondrites. CAIs are up to 2 million years younger than chondrules and thus indicate that the solid material of the solar system had a composition which was not only heterogeneous but also evolved over time (Norton, 1994).Lastly, what if the Earth accreted partly from differentiated planetesimals, as some recent studies have suggested (Halliday, 2006)? If so, then one cannot fairly assume a homogeneous, chondritic composition for the early Earth.
However, in the following sections a chondritic Earth will be assumed as this is the best estimate we have, currently, of early Earth’s composition. One must realize that this assumption of a chondritic early Earth is not without the concerns mentioned above.
The Accretion of the Earth:
Starting around 4.567 billion years ago (the age of chondrites), the Earth began accreting from planetesimals that most likely had a composition similar to CI carbonaceous chondrites (Taylor and Norman, 1990). Computer models suggest that most of Earth’s mass was accreted within ten million years of the formation of the solar system, and there was significant accretion up until about 100 million years after the formation of the solar system (Canup and Agnor, 2000). Actually, Earth is still accreting as 10^6-10^7 kg of meteorites are added to Earth each year (Norman, 1994). However, compared to the Earth’s current mass of 5.9736 * 10^24, this yearly accretion from meteorites is very small. However, in the early solar system meteorites played a significant role in Earth’s accretion as asteroids and planetesimals were much more abundant and had not yet settled into the current asteroid belts. Thus, collisions between these bodies were common in the early solar system. Also, Earth was not always so large– the planet grew over time from the accretion of meteorites.
After the solar nebula was formed, there were four main stages that led to the accretion of the Earth (e.g. Chambers, 2004; Boss, 1990):
1. The settling of circumstellar dust to the mid-plane of the protoplanetary disk. This stage occurred in thousands of years after the formation of the solar system.
2. The growth of planetesimals ~1 km in diamater through gravitational collapse of dense regions of the early solar nebula and through collisions of smaller particles.
3. Oligarchic (runaway) growth producing planetesimals ~1000 km in diamater. Note that this runaway growth can produce up to Mars-sized planetesimals but cannot produce an Earth-sized planet. Stages 2 and 3 together occur within a few hundred thousand years of the formation of the solar system.
4. Collisions between large, moon to Mars-sized planetesimals that lead to the formation of larger (i.e. Earth-sized) planets. The timing of this final stage of Earth accretion is debated, but Earth accretion likely occurred within 100 million years of the formation of the solar system, with most accretion occurring by 10 or even as little as 5 million years (e.g. Chambers, 2004; Hayashi et al., 1985).
There is significant evidence in the solar system that the final accretion of the Earth involved collisions between large impactors. First, asteroids are likely leftover planetesimals that were never accreted into a larger planet. The large asteroid belt between Mars and Jupiter likely formed because the strong gravitational pull of Jupiter kept these asteroids from accreting with the terrestrial planets (Norton, 1994). The presence of impact craters on planets and moons, especially on our own moon, also suggests that there were significant numbers of collisions between planetesimals in the early solar system. Modern-day asteroids still crash into planets, such as the famous impacts on Jupiter in 1994 when the Shoemaker-Levy 9 comet collided with the planet, and there were many more asteroids in the early solar system. Also, the offsets in the tilts of planets relative to their axes of rotation suggests that late in their accretion planets were hit by large impactors. That this offset exists supports a few collisions between large planetesimals late in planetary accretion. The affect of many small planetesimals colliding with accreting planets would not have an overall affect on the tilt of the planets as the effects of all the small impactors hitting the planet from all different directions would tend to cancel out (Taylor and Norman, 1990).
Part IV: It’s Getting Hot in Here, Differentiation, and Core Formation
Heating of the Early Earth:
Before moving on to a discussion of the differentiation of the accreted Earth, an understanding of the various processes that lead to the heating of the Earth is important. The early Earth was a very hot place and, indeed, the Earth is still fairly hot; plate tectonics and volcanism on Earth are largely driven by convection in Earth’s hot mantle.
There are four main sources of Earth’s heat (e.g. Lutgens and Tarbuck, 2003):
1. Potential (gravitational) energy that was converted to thermal energy during the accretion of the Earth.
2. The kinetic energy of the various impactors, which hit and heated the early Earth.
3. Differentiation of planetary layers. Again, gravitational potential energy was converted to thermal energy.
4. Energy released by the decay of both short-lived and long-lived isotopes. In particular, decay of short-lived isotopes such as 26Al may have contributed significantly to the heating of the early Earth and other planetary bodies.
One should also note that the presence of an early proto-atmosphere, perhaps comprised of gases captured from the solar nebula, may have helped blanket the Earth and prevented loss of heat through irradiation to space as the Earth was accreting. Later, atmospheres composed of volatiles released from accreting impactors may similarly have helped blanket Earth and retain heat.
Overall, the various sources of heat in the early Earth suggest that the early Earth was either partially or totally molten as the planet was accreting. Also, heating by impactors and the blanketing effect of a proto-atmosphere suggests that one or more magma oceans may have existed on Earth as the result of especially large impacts. In particular, a large magma ocean is believed to have formed after a Mars-sized impactor hit the proto-Earth in a collision that is believed to have produced Earth’s moon.
The Differentiation of the Earth:
After Earth accreted from the solar nebula, the Earth differentiated into several layers: a solid inner core, a liquid outer core, a solid but plastic mantle, and a thin shell of crust at the surface. The differentiation of the Earth’s core is one of the most significant events in Earth’s history and led to the concentration of iron as well as of siderophile and chalcophile elements at the center of the Earth. Most likely, the formation of the iron core was the first event in Earth’s differentiation. As will be described below, geophysical models and short-lived isotope systems, especially 182Hf-182W, can be used to place constraints on when and how core formation occurred in the Earth. Very likely, a deep, long-lived magma ocean– which formed when a giant impactor hit the Earth and created the moon– played a role in Earth’s core formation. After core formation, the upper Earth then differentiated into an enriched upper crust and a depleted mantle, which is the source material for modern oceanic crust. Again, geophysical models as well as isotopic systems and analysis of Rare Earth Elements (REEs) can provide constraints on how and when Earth’s upper mantle differentiated. Additionally, recent evidence from short-lived 142Nd isotopes suggests that there may have been an early protocrust, which was subducted deep in Earth’s mantle and has never been sampled.
Several models have been put forth to explain the formation of Earth’s iron core (see Figure 10). In early studies the core was modeled to have formed through a one-stage segregation of metal from a solid, fully-accreted Earth (Halliday, 2006). However, this simple model is probably unrealistic. More recent research has modeled Earth’s core forming gradually over millions of years and has examined the possible role of a magma ocean in the differentiation of the core. Additonally, some researchers have developed models of core formation in which Earth accreted (at least partially) from already-differentiated planetesimals (Halliday, 2006). If Earth was formed through giant impacts between differentiated planetesimals, the cores of these planetesimals may have merged to form the Earth’s core (Halliday, 2006).
Based on constraints from short-lived isotopic systems, Earth’s core is believed to have formed very early in Earth’s history. Most of Earth’s core was probably formed by about 10 million years after the formation of the solar system, and the final stage of core formation likely occurred in a magma ocean which resulted from a giant, moon-forming impact about 30 million years after the formation of the solar system (Jacobsen, 2005). Short-lived isotope systems have come into use recently as mass spectrometry has become sensitive enough to detect very small variations in these systems. Short-lived isotope systems allow geochemists to constrain the timing of core formation and other events in Earth’s history. The short-lived isotope systems 182Hf-182W (1/2-life: 9 million years) and 129I-129Xe (1/2-life: 16 million years) as well as the longer-lived U-Th-Pb system are all used to try to constrain when core formation occurred.
The 182Hf-182W system is the most straightforward system and has shed the most light on the formation of the core. Again, 182Hf has a 1/2-life of 9 million years, which means that variations created by the decay of this element cannot be detected after ~60 million years after the formation of the solar system (Lee and Halliday, 1995). Importantly, both Hf and W are refractory, which means that their ratio should not have been significantly altered during the accretion of the Earth from the solar nebula (Lee and Halliday, 1995). However, Hf is strongly lithophile while W is strongly chalcophile. This means that the Hf/W ratio would be greatly affected by core formation as W would be incorporated into the core while Hf would have remained concentrated in the upper, silicate Earth. The 182W/184W ratio of the Earth is slightly above chondritic (Jacobsen, 2005). This means that Earth’s core formation must have occurred before ~60 million years, after which time it would not be possible to significantly alter the 182W/184W ratio. By modeling the development of the observed 182W/184W ratios on Earth and the moon, scientists think that Earth’s core formed very early in Earth’s history and that the moon-forming impact occurred around 30 million years after the formation of the solar system (Jacobsen, 2005).
Part V: The Moon, the Magma Ocean, and the Mantle
The Moon and the Magma Ocean:
Since an early magma ocean may have played a key role in core formation, theories about the presence of a large magma ocean on the early Earth are important to understand. Many planetary geologists believe that about 30 million years after the formation of the solar system, a Mars-sized planetesimal collided with a mostly-accreted (~90%) Earth (Abe, 1996; Solomatov, 2000). The impact melted and vaporized a large portion of the upper Earth, and a magma ocean was formed and distributed over Earth’s surface. Additionally, the moon was created when a piece of the Earth was ejected into space by the impactor. Likely, the moon formed from a mixture of Earth and impactor material (Jacobsen, 2005). The magma ocean created by this impact may have taken as long as 100 million years to cool (Solomatov, 2000). A thick proto-atmosphere of either captured solar nebula gases or volatiles released from the Earth and the impactor as a result of the collision may have played a key role in blanketing the magma ocean, allowing radiation to be lost to space more slowly and the ocean to crystallize over a longer time (Abe and Matsui, 1986; Abe, 1997).
The exact depth and extent of this magma ocean is debated, but likely the magma ocean was fairly deep, potentially as deep as the modern core-mantle boundary, and crystallized in two stages (Abe, 1997; Solomatov, 2000). The deep magma ocean likely crystallized in two stages at two different rates (Solomatov, 2000). Interestingly, the magma ocean is believed to have crystallized from the bottom-up. This may seem counterintuitive at first as the Earth is hotter at depth. However, there is also higher pressure at depth. Figure 11 is a pressure and temperature diagram which shows the solidus and liquidus for orthopyroxene, a common mantle mineral, and several mantle adiabats. Because the mantle and, presumably, the ancient magma ocean are convecting, material is assumed to rise quickly without transferring heat– that is, the material rises adiabatically, cooling slightly as pressure decreases but remaining hotter, overall, than the upper mantle/magma ocean. Because the mantle adiabats in the magma ocean cross the solidus of mantle minerals at high pressures (i.e. deeper in the Earth), the magma ocean would have crystallized from the bottom up (Solomatov, 2000). Likely, the bottom part of the magma ocean crystallized in a short time period of about 1000 years (Solomatov, 2000).
The upper part of the magma ocean, the part above ~28 GPa, crystallized much more slowly over a time period of about 100 million years (Solomatov, 2000). Again, a blanketing proto-atmosphere– as well as subsequent impacts which re-melted the upper part of the magma ocean– enabled the magma ocean to exist for this long time period. Figure 12 illustrates the proposed two-layer model of the magma ocean.
If a deep magma ocean did exist on Earth for this long period, then this magma ocean may have played a significant role in core formation, which is believed to have occurred around the same time. Some researchers believe that an iron-rich layer settled to the bottom of the molten part of the magma ocean (Solomatov, 2000). Researchers have noted that it is difficult for metal-rich iron and sulfide melts to migrate through a static, solid silicate matrix (Rushmer et al., 2000). Immiscible iron and sulfide melts are able to travel through a silicate liquid much more easily (Rushmer et al., 2000). Thus, modeling the segregation of the iron core in a magma ocean is much easier than trying to model the segregation of the iron core in a solid Earth.
If an iron-rich layer did form at the base of a molten magma ocean, then Rayleigh-Taylor instabilities would have developed as this iron-rich layer is much denser than the lower magma ocean. Eventually, this instability would have caused the iron melt to travel downwards. The melt may have done this as large diapirs (see Figures 13 and 14) or by percolation in-between crystal grains (Figure 15). In a way, the migration of iron-rich material downward in the Earth can be thought of as reverse volcanism. Volcanic melts are believed to travel upwards through diapirs and percolation in-between mantle crystals. The core may have formed through similar mechanisms, just reversed– denser melts just migrated downward instead of lighter melts migrating upwards.
There is still much investigate about the characteristics of the magma ocean and how this ocean may have affected the differentiation of the Earth. Potentially, there was more than one magma ocean. There may have been a large magma ocean associated with the formation of the moon and several smaller magma oceans associated with other impacts (Abe, 1997). Thus, scientists must learn about the nature of different types of magma oceans: deep verses shallow, short-lived verses long-lived, and soft (high melt fraction, low viscosity) verses hard (low melt fraction, high viscosity) (Abe, 1997). Scientists also need to learn more about blanketing atmospheres associated with magma oceans, the nature of viscosity and convection in a magma ocean, and the role that crystal size and crystal kinetics in the crystallization of a magma ocean (Abe and Matsui, 1986; Solomatov and Stevenson, 1993).
Differentiation of the Upper Earth:
Since the focus of this paper is the origin and development of the early Earth, only a short summary of the later differentiation of the upper Earth will be presented here. In brief, then: after the formation of the core, the upper Earth differentiated into enriched, continental crust and a depleted mantle, which is the source for the oceanic crust (Best, 2003). Enriched means that the crust contains high abundances of elements, such as rare Earth elements, which are incompatible in mantle minerals and preferentially go into melts. Thus, the continental crust is believed to have formed first from the upper Earth and was followed by the formation of the denser oceanic crust. The abundances and patterns of incompatible elements, especially the light rare Earth elements, in oceanic crust suggests that the source for this crust was depleted by the formation of the continental crust (Best, 2006). Also, enriched means that the certain isotopic ratios are high while others are low. For instance, the continental crust has low 143Nd/144Nd and high 87Sr/86Sr ratios relative to oceanic crust (Best, 2003; Dickin, 2006). This reverse isotopic trend can be explained because Sm (the parent of 143Nd) is more compatible than Nd while Rb (the parent of 87Sr) is less compatible than Sr.
There is some debate in geology as to the timescales over which the continental and oceanic crust formed. However, the oldest crustal rocks are the 4 billion year old Acasta gneiss in Canada, so the continental crust was at least partly formed by this time (Valley, 2006). The oldest minerals on Earth are 4.4 billion year old zircons from Australia, and the existence of zircons this old suggests that at least small amounts of Granitic proto-continent existed by that time (Valley, 2006). Determining when oceanic crust first formed on Earth is more difficult as dense oceanic crust is generally subducted and recycled, which means that most of Earth’s oceanic crust is younger than 200 million years. The world’s oldest ophiolite (a section of oceanic crust thrust up over continental crust, usually at a back-arc basin) is the ~3.8 billion year old Isua ophiolite in Greenland, so oceanic crust was formed by at least this time period (Furns et al., 2007).
In addition to the continental and oceanic crust observed on Earth today, recent study of short-lived 142Nd isotopes suggests that there may have been an early protocrust, perhaps formed as a result of the crystallization of the magma ocean (Caro et al., 2005; Boyet and Carlson, 2005). This protocrust is called upon to explain anomalies in the 142Nd/144Nd ratio in Archean rocks (Caro et al., 2005). Supposedly, this protocrust was subducted in the Earth and has never been sampled (Caro et al., 2005; Boyet and Carlson, 2005). However, no one has yet put forth a reasonable physical mechanism by which the subduction of this protocrust could have occurred. Potentially, the observed 142Nd/144Nd anomalies are result of flaws in geochemical models assuming a chondritic starting material for the Earth rather than from a mysterious protocrust that was subducted and has never been sampled.
Summary: The Origin of the Earth in a Nutshell
Our solar system evolved from the solar nebula, which was composed of stardust from extinct stars and thus rich in heavier elements relative to cosmic abundances. Likely triggered by a shockwave from a nearby exploding supernova, the solar nebula collapsed gravitationally to evolve into the solar system. The solar nebula heated up, began spinning faster, and formed into a disk. Eventually, a proto-sun formed at the dense, hot center of the young solar system. At the same time, gases and dust began to condense in the outer, cooler parts of the solar system. Heavier, more refractory elements condensed closer to the sun, forming the terrestrial planets, while hydrocarbons condensed further from the sun, forming large bodies which were able to capture gases from the solar nebula and develop into gas giant planets. At the furthest reaches of the solar system, icy planets formed from methane, water, and ammonia ice. The Earth is believed to have accreted from chondritic planetesimals about 4.567 billion years ago. Chondrite meteorites come from old, undifferentiated asteroids that have undergone very little alteration or metamorphism. Carbonaceous chondrites are rich in organic material and are the least altered and metamorphosed of the chondrites. Thus, carbonaceous chondrites are often used as the starting material for Earth in geophysical and geochemical models. Earth was mostly accreted by ~10 million years after the formation of the solar system, and there was probably significant accretion of the Earth up to ~100 million years. Towards the end of Earth’s accretion, impacts between large planetesimals may have played a key role Earth’s growth and development. In particular, an impact from a large, Mars-sized impactor ~30 million years after the formation of the solar system may have created the moon and a deep magma ocean on Earth. The formation of the core probably occurred gradually as Earth accreted. However, the final stage of core formation may have been aided by the descent of iron and sulfur rich melts through a molten, silicate magma ocean. Eventually, the magma ocean crystallized, and the upper Earth differentiated into an enriched, continental crust and a depleted mantle, the source for oceanic crust. Potentially, an early proto-crust may have existed early in Earth’s history.
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