Supervolcano: A Geologic History of the Jemez Area

Kent G. Budge

Twelve thousand centuries ago, long before the first humans appeared in what would someday be northern New Mexico, the stillness was shattered by a series of explosions dwarfing anything ever witnessed by modern man. A giant chamber filled with molten rock and gas had long been growing under the earth, and when this finally exploded to the surface, the results were catastrophic. A mixture of red-hot gas and volcanic ash poured across the earth and covered the ground for miles in every direction. Finer ash was deposited up to 500 kilometers (300 miles) downwind, and the smallest particles were lofted into the stratosphere, dimming the sun for months. The roof of the emptied chamber collapsed to produced a volcanic crater, or caldera, 19 km (12 miles) across and 1000 m (3500') deep.

The caldera can still be seen on satellite photographs today.

Valles Caldera from space

Remnants of the thick layers of volcanic ash deposited by the eruption form the characteristic finger mesas of the Los Alamos area.

Pueblo Canyon from Clinton P. Anderson Scenic Overlook
Clinton P. Anderson Scenic Overlook. As with most images at this site, you may click to get a full resolution version of this image. 35 52.391N 106 14.006W

But the eruption of 1.2 million years ago was not the first in the Jemez area. An even larger eruption had taken place 400,000 years earlier, and smaller eruptions have occurred in the area for at least the last 14 million years. The last of these smaller eruptions took place about 55,000 years ago. Further eruptions are almost certain to take place, though there are no indications that we are likely to see such an eruption in our lifetime. This long history of eruptions has made the Jemez a choice location for studying the geologic processes involved in volcanic activity.

In the pages that follow, I will tell the story of the Jemez Mountains, from the assembly of northern New Mexico 1.8 billion years ago through the ages to the present. You will see photographs of some of the rock outcroppings that have allowed geologists to piece together this story. I will provide relief maps showing important locations, generated from NASA terrain elevation data with locations of roads and cities from the OpenStreetMap database. Geographical coordinates are provided for the adventurous reader who wishes to visit these locations using a GPS navigator. The geographical coordinates are also linked to Google Maps, as a further aid to navigation and overall understanding.

I have written these pages for readers who have an interest in science, but are not necessarily geologists. Because I am writing for a scientifically-minded international audience, I use metric units, but I also include English units for the benefit of the American reader.

Table of Contents

Chapter 1. Geology of the Jemez Area

Chapter 2. The Basement

Chapter 3. The Paleozoic

Chapter 4. The Mesozoic

Chapter 5. The Early Tertiary

Chapter 6. The Birth of the Jemez

Chapter 7. Volcaniclastics and Erosion

Chapter 8. The Tschicoma Formation

Chapter 9. Basaltic Volcanism Resumes

Chapter 10. Supervolcano!

Chapter 11. Resurgence

Chapter 12. The recent pass

Chapter 13. Wrapping it all up

Appendix: For The Hardcore Geology Enthusiast

Glossary and Bibliography

Introduction to the Jemez Area


Panorama of the Jemez Mountains from the southeast. 35 52.975N 106 03.660W

The Jemez Mountains are located in northern New Mexico, some 90 km (55 miles) north of Albuquerque and 55 km (35 miles) west-northwest of Santa Fe. The largest city in the Jemez area is Los Alamos, located on the Pajarito Plateau that extends east of the mountains. Other population centers in the Jemez area, starting southeast of Los Alamos and working clockwise, include White Rock, Cochiti, Jemez Springs, Cuba, Coyote, Abiquiu, and Espanola. There are no large population centers in the heart of the range, most of which is either National Forest land, part of the Valles Caldera National Preserve, or part of Bandelier National Monument or Kashe-Katuwe Tent Rocks National Monument. Much of the land that does not belong to the federal government belongs to Native American tribes, and most tribal lands are not open to the public.

The Jemez Mountains are best known to geologists for the Valles Caldera, the depression in the Earth's crust formed by the catastrophic eruptions 1.2 and 1.6 million years ago. Compare the digital relief map below with the satellite photo earlier.

Relief map of the Jemez
Digital relief map of the Jemez area. ©2015 Kent G. Budge

You will see variations on this map throughout the pages that follow, highlighting important features of the Jemez as they are discussed.

In the beginning

Relief map of the Jemez
Artist's depiction of the formation of a solar system. NASA

Our story begins with the birth of the Earth itself.

The Earth formed about 4.55 billion years ago from dust and gas expelled by dying stars, mixed with primordial gas from the beginning of the universe. This dust and gas was very different in composition from the Earth of today, being mostly hydrogen and helium with less than 2% of all the other chemical elements combined. Of these, oxygen was the most abundant, followed by carbon, neon, iron, nitrogen, silicon, magnesium, and sulfur. Only oxygen, iron, silicon, and magnesium remain abundant in the Earth today. This reflects the chemistry of the different elements

The early Solar System was a violent place, blasted by intense radiation and violent shock waves from the restless young Sun. Only those elements that could take stable solid form were able to contribute to the formation of the Earth and other planets. Helium and neon form no solid compounds, and so they are found in the modern Earth in only the smallest traces. Hydrogen, carbon, and nitrogen formed gaseous molecules of water, carbon dioxide, and ammonia that became part of the Earth only as minor components of more stable solid minerals. Sulfur formed various sulfur oxides and metal sulfides; the oxides behaved like water, carbon dioxide, and ammonia, while the metal sulfides likely found their way into the earth's core early in its history, leaving only much smaller quantities in the Earth's crust. By contrast, oxygen combined with iron, silicon, and magnesium to produce very stable solid minerals that now make up the bulk of the Earth.

We have a good idea of the age of the Earth from several lines of evidence. One is the age of the Sun. Astronomers can calculate how long it would take for a star with the Sun's mass and composition to reach its present state, based on what we know about the rate at which it burns hydrogen to helium. These calculations can be checked against helioseismic measurements, which look at how the entire Sun vibrates to get clues about its interior state. The models can be further checked by comparison with nearby stars. These all give a reasonably consistent age for the Sun of about 4.5 billion years. If we assume that the Earth formed out of the same cloud of gas and dust as the Sun, then the Earth must be around 4.5 billion years old as well.

Radiometric dating

Another way to estimate the age of the Earth is from the decay of radioactive isotopes. The atoms of a particular chemical element, such as oxygen or uranium, all contain the same number of protons in their nuclei (eight for oxygen, 92 for uranium), but the number of neutrons can vary. Thus, while all oxygen atoms have eight protons in their nuclei, some oxygen nuclei contain nine or ten neutrons rather than the usual eight. Each of the possible number of neutrons in the atoms of a chemical element defines an isotope of the element, and we distinguish different isotopes by adding a prefix to the element symbol that gives the total number of neutrons and protons in the nucleus. Thus the most abundant isotope of oxygen is written as 16O, since its nuclei contain eight protons and eight neutrons.

Some isotopes are unstable, and randomly decay to a different isotope, either by emitting two protons and two neutrons in the form of a helium nucleus (alpha decay) or by transforming a neutron to a proton by emitting an electron and a neutrino (beta decay). Alpha decay is characteristic of very heavy nuclei, while beta decay is characteristic of nuclei that are overloaded with neutrons. For example, the most common isotope of uranium, 238U, is the heaviest isotope of any element found in nature in any quantity. A nucleus of 238U is very slightly unstable, so that out of a large number of 238U nuclei, half will experience alpha decay over a period of 4.47 billion years. This period of time is known as the half-life of the isotope.

When a 238U nucleus decays, it is transformed into a nucleus of an isotope of thorium, 234Th. This nucleus is badly overloaded with neutrons and rapidly beta decays, with a half-life of just over 24 days, to an isotope of protoactinium, 234Pa. This is even more unstable, and undergoes another beta decay with a half life of just over a minute to become 234U. And so on, down through a decay chain of unstable nuclei, none of which have a half life greater than a quarter of a million years, until a stable isotope of lead, 206Pb, is reached.

Zircon grains
Zircon grains from sand deposits. USGS

We know the half life of 238U with high precision. We can also accurately measure the abundances of 238U and 206Pb in a rock sample using a suitable mass spectrometer. If we know how much of these isotopes the rock started with, we can then calculate its age. A particularly good mineral for this purpose is zircon, ZrSiO4, which is present as small grains in many kinds of rock. When a grain of zircon forms, it easily incorporates traces of uranium into its structure, but not lead, which has the wrong chemical behavior. As a result, we can assume that almost all the lead present in a zircon grain was produced by decay of the uranium it started with. If the amount of 238U is equal to the amount of 206Pb, then the age of the zircon grain must be equal to the half-life of 238U, 4.47 billion years. Younger grains will have a corresponding lower ratio of 206Pb to 238U, from which their ages may be calculated.

Zircon is a highly durable mineral, being hard and chemically inert. The oldest zircon grains found on Earth, from the Jack Hills of western Australia, are about 4.40 billion years old. This only tells us that the Earth cannot be younger than this age, because the oldest zircon may have formed some time after the Earth formed. However, we can also measure the age of rocks brought back from the Moon by the Apollo astronauts, whose landing sites included some that were deliberately chosen by NASA for their apparent great age. The most accurate age measurements of all come from meteorites, some of which (known as carbonaceous chondrites) are virtually unchanged from the very beginning of the Solar System. It is from these meteorites that we get a precise age of 4.55 billion years for the Solar System and, presumably, the Earth.

The actual process of formation of the Earth involved dust grains merging to form progressively larger bodies, called planetesimals, that merged through further collisions to form the planets. This process was mostly complete by 10 million years after the birth of the Solar System.

The Earth's core and mantle form

The newborn Earth was extremely hot and molten, both from the heating caused by the impacts of the planetesimals that became part of the Earth, and because the young Earth had a fair amount of radioactive material in it from the dying stars that contributed the heavier elements in its makeup. Much of the iron in the young Earth remained uncombined with oxygen. Being heavy, this iron settled into the center of the Earth, taking with it much of the sulfur and heavy metals with which the Earth was born.

We know this through another clever line of reasoning based on isotope measurements. One of the isotopes of hafnium, 182Hf, decays to an isotope of tungsten, 182W, with a half-life of just nine million years. Tungsten dissolves readily in liquid iron, but hafnium does not. Thus, any 182W produced before the earth's core formed would have been carried into the core along with the iron, while any 182W produced after the core formed would remain in the Earth's outer layers. The amount of 182W relative to other isotopes of tungsten in terrestrial samples is higher than the ratio in meteorites, showing that the core formed before all the 182Hf could decay into tungsten.

Large planetesimal collision
Artist's depiction of large planetesimal collision, similar to the Theia event. NASA

The same approach, but using the 235U - 207Pb isotope pair, gives an age for the core of about 80 million years. This is significantly longer than the estimate of 30 million years from radioactive hafnium. The likely explanation is that the Earth collided with a Mars-sized object, which planetary scientists have named Theia, very early in its history. This also explains the origin of the Moon, which is thought to have formed from debris left orbiting the Earth after the collision. This colossal impact reset the uranium clock for core formation, but not the hafnium clock, which was already run down by that time.

Another line of evidence is the abundance of chemical elements in the mantle compared with carbonaceous chondrite meteorites. I have already mentioned that tungsten dissolves readily in molten iron. A number of other elements, which geologists classify as siderophiles, have a similar affinity for iron. Their pattern of depletion from the mantle tells us that the iron core must have formed very early and under conditions of high pressure, most likely in the depths of an ocean of molten rock (which geologists call magma). In fact, some planetary scientists believe that iron had already begun separating into cores in the planetesimals that merged to form the Earth.

Thus, within about 30 million years, the Earth had a core of iron enriched with sulfur and heavy metals, surrounded by a mantle composed mostly of silica (silicon dioxide, SiO2) and metal oxides. It was at about this time that the early Earth was struck a glancing blow by Theia. Most of the core of Theia became part of the Earth, explaining why the Earth has a relatively large iron core, while much of the mantle of Theia went into orbit around the Earth and eventually coalesced into the Moon, which is relatively poor in iron. The tremendous energy of the collision melted most of the Earth's mantle, turning it back into an ocean of magma.

Large computer simulations of the impact suggest that parts of the magma ocean may have reached temperatures as high as 7000 K (12,600 F), briefly giving the Earth at atmosphere of vaporized rock. However, this would have cooled very quickly. As the Earth continued to cool, convection currents in its young mantle helped heat escape, and the mantle rapidly solidified, probably in less than 10 million years.

Convection is still an important process in the Earth today, so it's worth taking a moment to explain. Almost all materials contract as they cool. (Water freezing to ice is a notable exception.) When a hot body of liquid is allowed to cool, its exposed surface cools first. The cooler surface liquid contracts and becomes denser than the hotter liquid beneath. This causes it to sink, and hotter material from beneath rises to take its place. This is the process of convection, and it is an important mechanism for transporting heat energy. 

Calculation of convection in the mantle
Computer simulation of convection in the Earth's mantle. Via Wikimedia Commons

In a viscous fluid, or in a fluid that is cooling very slowly, the rising and sinking currents tend to organize themselves into large cells, like those in the simulation shown above. Magma is fairly viscous, and the rate of cooling is very slow for a body as large as the Earth. Hence, convection within the Earth may generally be expected to organize itself into large convective cells.

The oceans and continents form

Like the core, the solid outer crust of the Earth formed very early in its history. I mentioned earlier that the oldest known individual zircon grains are about 4.40 billion years old, showing that the first solid rock must have formed within about 150 million years of the Earth's formation. Most likely a solid crust formed very shortly after the impact that created the Moon, but all traces of the earliest crust have been destroyed, either by later impact events, by erosion, or by other geologic processes.

This early period of the Earth's history, from which no solid rock survives, has been named the Hadean Eon by geologists. The name alludes both to the hellish conditions that prevailed on the young Earth and to the hidden nature of this time period. The oldest known whole rock, found at the Acasta River in the Northwest Territories of Canada, is about 4.055 billion years old, and geologist have placed the end of the Hadean Eon at 4 billion years ago. The Hadean was followed by the Archean Eon, the first from which we have samples of solid rock.

Magnetite schist from the Tusas Mountains
Banded iron formation from the Tusas Mountains. Banded iron formation is not found in the Jemez. 36 38.331N 106 8.072W

Geologists are reasonably certain that there was liquid water on the Earth's surface by 3.8 billion years ago. This is the age of the oldest known example of a banded iron formation, at Isua, Greenland. Banded iron formation is a kind of rock consisting of thin, alternating layers of silica and iron oxides, and it forms by chemical precipitation in liquid water. However, the oldest oceans may have been closer to 4.2 billion years old, based on oxygen isotope ratios in the oldest zircon grains. The ratio of 18O to 16O increases in rock that reacts chemically with ocean water, because the slightly lighter and more mobile water molecules containing 16O are more likely to evaporate from ocean water or, when ocean water reacts with rock, to remain in solution. When  the 18O-enriched rock is recycled to magma in the lower crust and this magma forms new zircons, the zircons also show the enhanced 18O. The 18O/16O ratio in zircon grains increased significantly at around 4.2 billion years ago, and some geologists interpret this as the time the first large oceans formed.

These oceans would have been very different from the oceans of today. They formed under an atmosphere containing practically no free oxygen but thick with carbon dioxide, so the earliest oceans would have been anoxic and highly acidic. The ocean water would have rapidly leached sodium and other soluble ions out of the crust to become saline. Limestones and other carbonate rock of Archean age is quite uncommon, because they would have tended to dissolve in acidic ocean water.

At the beginning of the Archean, the earth's crust was almost entirely what we would now classify as oceanic crust, whether or not it was all covered with oceans. To understand what this means, we need to jump ahead of our story and look at the structure of the Earth as it is today. This will be a long digression, but it will lay the groundwork for much of the rest of this book.

Diagram of convection in the Earth's

Diagram of mantle convection. U.S. Geological Survey

We've already seen that the original materials of the Earth quickly separated into an iron core and a silicate mantle. Both were initially molten. The core has since partially solidified, starting at its center, to produce a solid inner core. The outer core is still liquid, and the slow motion of the convective currents in the outer core are responsible for the Earth's magnetic field.

The mantle is almost completely solid. However, there is a layer in the upper mantle, the asthenosphere, which contains scattered pockets of magma. Even in these pockets, the magma is a small percentage of the total volume, filling pore space in otherwise solid rock. The asthenosphere begins about 80 km (50 miles) below the surface, but has no well-defined lower boundary. Different geologists estimate the thickness of the asthenosphere as anywhere from 110 km (70 miles) to several hundred kilometers. Because the asthenosphere is close to its melting point, it is a zone of weakness. This gives it its name: Greek asthenḗs = 'weak'.

This peculiar distribution of liquid regions in the Earth is the result of the competition between increasing temperature and increasing pressure as one goes deeper into the earth. Rock under the enormous pressure found deep in the Earth melts at a higher temperature that it would at the surface. In the crust and outermost mantle, the increase in pressure is enough to keep the rock from melting in spite of the increase in temperature with depth. Only in the asthenosphere does the temperature get just enough ahead of the pressure for the rock to partially melt. Further down, pressure gets ahead again, and the rest of the mantle is solid rock. At the boundary of the outer core, there is an abrupt change in composition from iron and magnesium silicates with a very high melting point to nickel- and sulfur-enriched metallic iron with a somewhat lower melting point. This is enough for the iron to remain molten. Deeper in the core, the pressure again catches up with temperature, and the inner core is solid.

In spite of being mostly solid rock, the mantle is still experiencing very slow convection. This is possible because the hot rock is ductile. If you hit a piece of lead with a hammer, the lead tends to deform under the force of the blow rather than shatter. The same is true of many other metals. Even iron will deform under force if it is heated to a high temperature, a quality that is used to advantage both by traditional blacksmiths and in modern steel mills. At the Earth's surface, silicate minerals shatter rather than deform, but at the high pressure and temperature in the Earth's interior, silicates become ductile. This ductile rock is able to flow very slowly, over geologic time scales, at rates of a few centimeters a year.

The depth at which brittle rock becomes ductile is known as the brittle-ductile transition, and this is normally about 13 to 18 km (8 to 11 miles) down. However, the transition can vary locally, as is proven by the existence of very deep earthquakes in certain geographical locations, such as Tonga. Earthquakes take place only in rock that is still brittle.

Oceanic versus continental crust

The earth's outermost layer, the crust, is not uniform. Geologists realized this when they began making systematic seismic observations in the late 19th century.

An earthquake takes place when brittle rock under stress suddenly breaks to relieve the stress. This disturbance sends elastic waves through the surrounding rock, which can be detected at great distances with sensitive seismographs. These waves move at a velocity that is determined by the nature of the rock they travel through as well as the nature of the wave itself.

Earthquake wave types
U.S. Geological Survey

The fastest waves are P waves, or pressure waves, which are similar to sound waves. A P wave is a zone of compression that moves through the material. P waves are capable of passing through almost any kind of material, including liquids and gases, and so they are observed passing through every part of the Earth's interior. An S wave, or shear wave, is a zone of shear, in which the material is pulled to one side before snapping back. Think of snapping a stretched spring to one side, and watching a kink propagate down the spring. S waves propagate only through solid materials, and they propagate more slowly than P waves. A skilled seismologist can determine how far away an earthquake took place by comparing the arrival times of P waves and S waves. There are also various kinds of surface waves, none of which penetrate deeply into the earth and all of which travel more slowly than either P or S waves. 

P waves travel more quickly through less compressible materials. As one goes deeper into the earth, the rock is under greater pressure, which means that it is already highly compressed and resists further compression. Thus, the velocity of P waves tends to increase smoothly with depth. However, seismologists have found that there are certain depths at which the velocity changes more abruptly. These are called seismic discontinuities. These discontinuities are interpreted as depths at which the nature of the rock changes. Such changes can be the result of a change in chemical composition, such as the change from silicates to iron at the core-mantle boundary, or they can be the result of a phase change in the rock, in which the composition is the same but the rock changes from one crystal structure to another that is more stable at the higher pressure.

The first seismic discontinuity to be identified was discovered by the Croatian seismologist Andrija Mohorovičić in 1909. He found that seismic wave velocities increased sharply a few kilometers beneath the surface. The Mohorovičić discontinuity (which most geologists refer to simply as the Moho) is now used to define the boundary between the upper mantle and the crust. It is believed to be the maximum depth at which a mineral called feldspar is an important component of the rock. I'll have more to say about feldspar later in this book.

As geologists mapped the Moho, they found that the thickness of the crust was much greater under the continents than in the ocean basins. This was not surprising. Geologists had already concluded from theoretical principles that mountain ranges must be held up by low-density roots, just as the tip of an iceberg is held above the surface of the ocean by the buoyancy of the rest of the iceberg lying beneath the ocean's surface. The existence of these mountain roots was confirmed by surveyors in India in the mid-19th century, who discovered that their plumb bobs pointed slightly away from the Himalayas. The low-density roots of the mountains distorted the Earth's gravitational field. Geologists now perform gravimetric measurements using extremely sensitive gravimeters to get clues about what lies underground.

What is striking is how different continental crust is from oceanic crust, and how abrupt the transition from one to the other is. Continental crust forms almost all the dry land on Earth, plus the continental shelves, which are the areas of relatively shallow water around the margins of the continents. Some distance offshore, the continental shelves begin to slope steeply downwards to the deep ocean floor, which is composed of oceanic crust. 

Diagram of continental shelf
U.S. Navy. Via Wikimedia Commons

Oceanic crust is thin (less than 10 km or 6 miles thick) and dense (about 3.0 g/cm3) compared with continental crust, which is 25 to 65 km (16 to 40 miles) thick with a density of  about 2.7 g/cm3.  Both types of crust are enriched in silica, highly enriched in aluminum, sodium, potassium, and calcium, and depleted in magnesium compared with the underlying mantle, which has a density of about 3.4 g/cm3.  However, oceanic crust is less enriched with these elements than continental crust.

The difference between oceanic and continental crust reflects different origins and has important consequences. 

Plate Tectonics

        seafloor crust age
Chart of sea floor age. Click to enlarge. NOAO

As early as 1850, the great American oceanographer, Matthew Fontaine Maury, suspected the existence of a underwater mountain range in the central Atlantic. Tantalizing evidence for such a range was uncovered by the British HMS Challenger during its famous 1872 expedition, but it was not until the 1950s that a team of oceanographers from Columbia University, led by Marie Harp and Bruce Heezen, fully mapped the great underwater mountain chain running from north to south through the center of the Atlantic Ocean. Other researchers found that this chain, dubbed the Mid-Atlantic Ridge, continued through the Indian Ocean and the far southern Pacific Ocean before turning north as the East Pacific Rise and continuing to the west coast of North America. This mid-ocean ridge has a deep valley along most of its crest, where there is volcanic activity and frequent earthquakes. Geologists also found that the rock is very young close to the central valley and grows progressively older further away, in either direction. The implications were clear: The central valley of the mid-ocean ridge is a rift valley where the crust is being pulled apart, and volcanic activity creates new crust to fill the resulting gap.

The discovery of the mid-ocean ridges marked a turning point in the acceptance of the theory of plate tectonics. According to this theory, the mantle is in slow convective motion. Hot, ductile mantle rock continually creeps upwards beneath the mid-ocean ridge. As the rock approaches the surface, the decrease in pressure allows it to partially melt. The magma so produced erupts along the rift valley of the mid-ocean ridge to form new oceanic crust, which then moves away from the ridge.

Map of tectonic plates
Map of tectonic plates of the modern Earth. U.S. Geological Survey

The crust forms the upper portion of the lithosphere, which is the rigid outer shell of the earth. The uppermost layer of the mantle forms the lower portion of the lithosphere. Beneath the lithosphere is the ductile asthenosphere. The lithosphere is broken into large plates that are displaced by the convection currents in the underlying mantle.  A single plate may include both oceanic and continental crust. For example, the North American plate includes much of the floor of the western North Atlantic Ocean as well as the continent of North America.

The distinction between oceanic and continental lithosphere is crucial to the behavior of plates. A region of lithosphere surfaced with light, thick continental crust is too buoyant to sink into the mantle. As a result, continental lithosphere may be torn in half (rifted) or merged with another block of continental lithosphere (sutured) but it is almost never pulled into the mantle and destroyed. On the other hand, a region of lithosphere surfaced with thin, heavier oceanic crust is nearly as dense as the underlying asthenosphere. When it is freshly formed and still hot, oceanic lithosphere is is buoyant enough to resist sinking into the mantle, but as it slowly cools and contracts, it reaches a point where it is no longer buoyant.

What goes up must come down, and the convective flow of hot mantle rock upwards beneath mid-ocean ridges must be matched by a downward flow of cold mantle rock somewhere else. In the modern earth, this takes place where plates collide.

When cold oceanic lithosphere meets buoyant continental lithosphere, two things can happen. If the oceanic and continental lithosphere are able to move together as a single plate, like the continental and oceanic portions of the North American Plate, then the continental margin is described as a passive margin. The continental shelf ends in a continental slope that drops smoothly to the ocean floor. There is no particular seismic or volcanic activity along a passive margin.

Diagram of destructive margin
Wikimedia Commons

On the other hand, if the continental and oceanic lithosphere are driven together, the thick, light continental lithosphere rides up on the thinner and heavier oceanic lithosphere. The oceanic lithosphere is driven downwards, and because it lacks buoyancy, it begins to sink into the mantle. This process is called subduction. The point where the oceanic lithosphere begins to sink into the mantle is marked by a deep oceanic trench, and ocean floor maps show such trenches along many continental margins. In particular, subduction is taking place along most of the continental margins of the Pacific Ocean and south of Indonesia in the Indian Ocean. These are described as destructive margins, and how!

The oceanic lithosphere does not go gentle into that good night. The motion of oceanic lithosphere under continental lithosphere produces the most powerful earthquakes known, as the plates periodically lock together, stress builds, and then the plates break loose over an area that can extend for hundreds of kilometers. The damage is compounded by tsunamis produced by the sudden movement of the ocean floor next to the trench. Historical examples of such megathrust earthquakes include the great Chilean earthquake of 1960, the most powerful ever recorded; the Alaskan "Good Friday" earthquake of 1964; the Indonesian "Boxing Day" earthquake of 2004, which produced a tsunami that killed nearly a quarter of a million people; and the Japanese Tohoku earthquake of 2011 that produced a tsunami that killed as many as 18,000 people and wrecked the nuclear power station at Fukushima, topping tsunami barriers 14m (45 feet) high.

Wadati-Benioff zone of the Kurile

Plot of earthquakes along the Wadati-Benioff zone of the Kurile Islands. USGS.

The descending slab remains relatively cold and brittle for some time. Such slabs are responsible for the deepest earthquakes, which occur at depths of up to 600 km (370 miles). In 1949, seismologist Hugo Benioff noted that the focuses of earthquakes in the Tonga area of the South Pacific and on the western margin of South America were located on planes dipping at an angle of about 45 degrees into the earth. Japanese seismologist Wadati Kiyoo made the same observation at about the same time, and these zones of deep earthquakes, corresponding to subducting lithosphere, are now known as Wadati-Benioff zones. The existence of these zones was important evidence supporting the theory of plate tectonics.

What is the force that drives the subducting plate into the subduction zone? Most likely it is the weight of the subducting plate itself. This produces slab pull that acts all the way back to the mid-ocean ridge. The asthenosphere on either side of the subducting plate is also pulled down with it, like the suction produced by a sinking ship, and this suction helps draw the overriding plate over the subduction zone.

Thus oceanic crust is constantly being created and destroyed, as it moves like a conveyor belt from its birth at mid-ocean ridges to its death at destructive margins. No oceanic crust on Earth has an age greater than about 270 million years, which is a mere 6% of the age of the Earth. On the other hand, portions of the continental shields are as old as 4 billion years, or 88% of the age of the earth, because continental crust does not subduct. The picture we get from plate tectonics is of ancient continents that are continually jostled around by the movement of young oceanic crust, which in turn is driven by mantle convection.

Obviously subduction must have worked differently before there was any continental crust. We have no geologic evidence for how subduction worked in the Hadean, though some planetary scientists have suggested it resembled the hot spot volcanism now seen on one of Jupiter's moons, Io. However, by the beginning of the Archean Eon, subduction resembled what we see today in oceanic island arcs. Examples of such arcs can be found in the western Pacific and southern Atlantic. Here the subducting oceanic lithosphere is being overridden by a different oceanic lithosphere plate, rather than a continental lithosphere plate.

Destructive margins usually have a line of volcanoes (a volcanic arc) along the overriding side of the trench. The oceanic crust descending into the mantle contains a great deal of water, and this boils off when the subducting slab is heated at depth. Above the slab, on the overriding side, there is a mantle wedge into which the water rises. Water lowers the melting point of silicate minerals, and the more silica, the greater the decrease in the melting point. This is because water attacks the oxygen bonds between silicon atoms.

When a liquid solidifies, its atoms usually arrange themselves into regular lattices called crystals. This minimizes the chemical energy of the substance, by allowing each atom to chemically bond with as many other atoms as possible. (Chemical bonding lowers the energy of a group of atoms, which is the foundation of the science of chemistry.) How many bonds are possible is determined by the size of each atom, and by the arrangement of its electrons around its nucleus, which is determined by the rules of quantum mechanics. The specifics are different for each chemical element, and each stable arrangement of atoms into crystals is called a mineral. For example, the most common kind of crystal that forms from pure silica is the mineral, quartz.

Quartz sample

A sample of quartz from the Picuris Mountains. Quartz of this purity is rare in the Jemez. 36 11.997N 105 50.099N

Silicon atoms bond quite strongly to oxygen, and each silicon atom bonds to four oxygen atoms.

Silica tetrahedron

An isolated silica tetrahedon. Silicon is grey and oxygen is red. The atom sizes are reduced to better show the bonds.

Each oxygen atom, in turn, bonds to two silicon atoms.

Silica tetrahedra joined at corners

A pair of silica tetrahedra jonied at their corners

In minerals like quartz, the structure can be thought of as an arrangement of tetrahedra, each with a silicon atom at its center:and with each tetrahedron joined to a neighboring tetrahedron at each of its corners through a shared oxygen atom.  This produces a strong three-dimensional network of interlocked silica tetrahedra, and so it is unsurprising that quartz is the hardest common mineral.

The reaction of water with silica can be visualized as:

Reaction of water with silica

This breaks down the silica network and thereby lowers the melting point of the silica. The consequence of this is that, whenever a wet slab of crust sinks into the mantle, the water that boils out of the subducting slab and rises towards the surface causes partial melting of the overlying mantle wedge. This leads to the creation of new continental crust.

Partial Melting

The key is that the melting is partial. Mantle rock consists of a mixture of different minerals, all with different melting points. Such a mixture of different minerals begins melting at a lower temperature than any of the pure minerals alone. This is similar to the way that salt sprinkled on ice lowers its melting point.  The temperature at which a mixture of minerals first begins to melt is called the solidus, since below this temperature the mixture is completely solid, while above this temperature the mixture is at least partially melted.

The magma produced at the solidus has a very well-defined composition, called the eutectic. This is the composition that minimizes the chemical energy of the magma plus solid minerals. The rock will continue to produce magma of this eutectic composition as heat is added, until one of the solid minerals that contributes to the eutectic is fully melted. The melting process will then shift towards a new eutectic at a higher temperature determined by the remaining solid minerals.

But the process almost never reaches the point where all the solid minerals have melted, because there is very rarely enough heat available to fully melt the rock. Typically less than 30% of a mantle source rock will be melted before the supply of heat energy is exhausted. Because the eutectic can (and usually does) have a different composition than the solid source rock from which it came, the effect of partial melting is to produce magma with a composition different from the source rock, whose own composition is also changed by the extraction of the magma.

When typical mantle rock is partially melted, the magma produced has a composition of about 50% silica, versus 40% for the source rock. Magnesium is very reluctant to enter the eutectic, so the magma contains perhaps 16% magnesium oxide versus about 50% for the source rock. Calcium and aluminum are enriched about fourfold in the magma, while the alkaline metals, sodium and potassium, are greatly enriched, by a factor of 10 to 30.  This magma has roughly the same composition as the common volcanic rock, basalt.

Lobatto Formation basalt from
        Clara Peak
Basalt from the Cerros del Rio. 35 49.285N 106 11.069W

The remaining solid mantle rock is depleted in the elements that are enriched in the magma. Geologists believe that a rare type of rock sometimes found as inclusions in lava flows, called harzburgite, is depleted mantle rock caught up in the magma. 

The magma is lower in density than the source rock, so its buoyancy drives it towards the surface, where it solidifies into rock that is lower in density than the mantle from which it came. Most of the oceanic crust of the Earth is composed of basalt or rocks similar to basalt. Basalt is nevertheless a relatively dense crustal rock, and when oceanic crust made of basalt cools sufficiently, it is capable of sinking back into the mantle by subduction.

Magma Differentiation

Continental crust is composed of rock which, on average, is significantly less dense than basalt. It follows that the formation of continental crust must involve some further process that changes the composition of the original (primitive) basalt magma. This process is magma differentiation

As magma rises towards the surface, it begins to lose heat to the cooler solid rock that surrounds it (the country rock), As it cools, it begins to crystallize.

Just as source rock does not melt all at once, and the partial melt is different in composition from the original source rock, so magma does not usually solidify all at once. There is a temperature (the liquidus) at which the magma first begins to crystallize, and the first crystals that form usually do not have the same composition as the magma. Instead, the first crystals that form have whatever composition drives the remaining magma towards its eutectic. This is similar to the melting process, but in reverse. In fact, if the magma had never left its source region, the crystallization process would be precisely the melting process in reverse. It is not precisely the same because the magma has left behind the most refractory (highest-melting) components of the source rock and is now in a cooler, lower-pressure environment.

Bowen's reaction series
Bowen's reaction series. Via Wikimedia Commons

The process of crystallization of a magma was first studied by Norman Bowen in the early 1900s. Bowen melted various rock powders and observed which minerals formed first when these artificial magmas were slowly cooled. He found that there were two distinct series of minerals that crystallized out of a cooling magma. Bowen's discontinuous reaction series was a sequence of minerals that crystallized one at a time, but there was also a continuous reaction series of increasingly calcium-poor and sodium-rich plagioclase feldspar. We'll have more to say about most of these minerals later in the book. For now, we note that the discontinuous series removes magnesium rapidly and iron more slowly from the magma, leaving a liquid that is increasingly enriched in silica, aluminum, and the alkaline metals. The continuous series, in turn, removes calcium from the magma.

Minerals such as olivine and pyroxene are denser than basalt magma. Given enough time, crystals of these minerals will slowly settle out of the magma. Plagioclase is less dense and often remains suspended in the magma, or even floats to the top of a dense magma body. As a result, magma that has partially crystallized is less dense than the original basalt magma and often contains suspended crystals of plagioclase. This differentiated magma has an intermediate silica content, about 57% to 63% by weight, and if it erupts to the surface, it will harden into a characteristic kind of rock called andesite. Andesite is poorer in calcium and magnesium as well as somewhat richer in silica than basalt. Whereas basalt is usually a fairly uniform dark rock, andesite typically contains large individual crystals (phenocrysts) of plagioclase and hornblende.


Andesite from the Paliza Canyon Formation, St. Peter's Dome. 35 45.799N 106 22.396W

The process of magma differentiation can continue past andesite to produce increasingly silica-rich magma. Magma with a silica content of 63% to around 70% produces a rock called dacite.

Dacite of the Pajarito Mountain Member, Tschichoma Formation, Camp May. 35 53.807N 106 23.832W

Magma whose silica content has reached 70% or more forms a rock called rhyolite

Rhyolite of Rabbit Mountain, Valle Toledo Member, Cerro Toledo Formation. 35 49.637N 106 28.075W

You may have noticed that, as the rock becomes increasingly rich in silica, it also becomes lighter in color. This is because magma differentiation removes iron from the magma as it enriches it in silica. This is a useful rule of thumb rather than a rigid rule; some dacites can be quite dark in color, as we'll see later in the book. However, almost all basalts are quite dark in color, and almost all rhyolites are light colored.

Basalt, andesite, dacite, and rhyolite are all examples of extrusive volcanic rocks. They are formed from magma that reaches the surface and then cools so rapidly that it does not have time for the liquid magma to form large crystals. Extrusive rocks are therefore very fine-grained, except for phenocrysts that had already formed in the magma before it reached the surface. It is common for magma that solidifies underground, before it can reach the surface, to have a much coarser texture, and these intrusive rocks are given names of their own. The intrusive counterpart of basalt is called gabbro (or, if somewhat finer in texture, diabase):


Diabase from the Lobato Formation, Los Cerros. 36 01.680N 106 15.247W

The intrusive counterpart of andesite is called diorite.

Diorite from the Abajo Mountains. Diorite is rare in the Jemez.

The intrusive counterpart of dacite is called granodiorite.

Maquinita Granodiorite

The intrusive counterpart of rhyolite is called granite.

Joaquin Granite
Joaquin Granite. 35 46.421N 106 47.519W

Igneous rocks low in silica are described as mafic, a word which was coined to reflect the high MAgnesium and Ferrous oxide content of these rocks. Likewise, the word felsic was coined to reflect the high content of FELdspar and quartz in high-silica igneous rocks. Andesite is described as an intermediate rock. There are many further refinements to this simple igneous rock classification scheme. For example, rocks midway between basalt and andesite are described as basaltic andesites, and rocks midway between dacite and rhyolite are described as rhyodacites. And since silica is only one component of igneous rock (albeit the most important) there are further classifications based on the abundances of the other oxides in the rock, particularly the akaline metals, sodium and potassium. Rather than bog you down further with a discussion, say, of the distinction between tholeiitic and alkalic basalt, or of the finer points of distinguishing a rhyodacite from a quartz latite, I'll introduce these terms (if necessary) when I show pictures or samples of particularly well-characterized rock formations. If you are really interested, the full classification scheme is found in the appendix.

A magma composed of pure silica would be extremely viscous, because the silicon and oxygen atoms keep trying to arrange themselves into silica tetraheda even in the liquid state. This produces localized chains and clumps of tetrahedra that impede flow. Metal atoms break up these chains and clumps and allow the melt to flow more freely, so low-silica mafic magmas are much less viscous than high-silica felsic magmas. Basaltic magma has about the viscosity of ketchup and flows relatively freely. Andesitic magma has enough silica to be about as viscous as smooth peanut butter, and so it flows significantly less freely. Dacitic magma is more viscous still (with a viscosity similar to Silly Putty) while rhyolitic magma has great difficulty flowing at all.

Subduction in oceanic island arcs produces a great deal of andesite, and this was likely the nucleus of the first continents. Further subduction produced increasingly large quantities of dacite and rhyolite and their intrusive counterparts, as the thickening crust meant that the magma took longer to reach the surface and had more time to differentiate. In some cases, the primitive basalt magma never penetrated the crust at all, underplating the crust instead. The hot basaltic magma at the base of the crust heated the rocks above, causing them to melt and rise to the surface as silica-rich magma. The net effect was to concentrate the iron and magnesium at the base of the crust, where it sometimes delaminated or came loose and sank back into the mantle. The crust became increasingly thick, rich in silica, and low in density. It became continental crust.

There is evidence that the bulk of the Earth's continental crust formed by 2.5 billion years ago. Once again, the necessary clues are provided by radioisotopes. The rare earth element, neodymium, is slightly more likely to enter the partial melt that produces crust than is samarium, and so the decay of 147Sm to 143Nd tracks the formation of crust. Decay of 87Rb to 87Sr works in a similar way, with rubidium being far more likely to enter a partial melt than strontium. Both isotope ratios suggest that formation of continental crust had peaked by about 2.5 billion years ago. The period from 4 billion years ago to 2.5 billion years ago has been named the Archean Eon by geologists.

You may have noticed that I haven't mentioned the Jemez Mountains in my story so far. This is because New Mexico did not exist until 1.8 billion years ago. In the next chapter, I'll tell the story of how New Mexico first came to be.

Next chapter: The basement

Copyright ©2014 Kent G. Budge. All rights reserved.