Geology of the Jemez Area, Chapter 2: The Basement

Table of contents here.

Amchitka Island
Amchitka Island. The Jemez area might have resembled this scene 1.7 billion years ago.

The Earth of 1.8 billion years ago was a very different world than the Earth of today. The atmosphere was thick with carbon dioxide and had less than 10% of its current abundance of oxygen. Though there were continents, they were wastelands barren of life, and even the oceans contained only primitive microorganisms. It was in this setting that northern New Mexico first came into existence.

The first chapter of this book began the story of the Jemez Mountains with the formation and early history of the Earth. In this chapter, we will look at the oldest geologic features of the Jemez area.

The Precambrian

Map of the Jemez highlighting Precambrian
Relief map of the Jemez with Precambrian outcroppings highlighted in red.

In the early days of scientific geology, geologists found that sedimentary rocks (rocks formed from sediments eroded from older rocks) often had distinctive collections of fossilized organisms in them. In most locations, the lowest sedimentary rock beds contained fossils of more primitive forms of life than the higher beds. Geologists worked out a regular progression from the most primitive fossils to fossils much like animals seen today. This allowed sedimentary beds found at widely separated locations, but containing similar fossils, to be correlated. Although the absolute age of the rocks could not yet be determined, the relative age could. Geologists worked out a time scale based on relative age and began giving names to each interval of geologic time. The three main eras in the fossil record were named the Paleozoic ("ancient life"), the Mesozoic ("middle life"), and Cenozoic ("new life"). Each era was further broken down into periods, such as the Cambrian Period at the beginning of the Paleozoic Era.

Geologists recognized that, in many places in the world, there were rock layers beneath the Cambrian Period beds that contained no fossils. These Precambrian rocks, as they are often still called today, are a mess. They tend to be coarsely crystalline rocks, either intrusive volcanic rocks or metamorphic rocks (rocks recrystallized under great heat and pressure.) They are often highly deformed and fractured. Precambrian sedimentary beds could not be correlated because they contained no fossils. Thus this Precambrian "basement" was all but indecipherable.

The discovery of radioactivity led to the invention of radiometric dating of rocks in 1907 by American geologist Bertram Boltwood. Geologists were finally able to assign absolute dates to the various periods in the geologic record. They discovered that the oldest Cambrian rocks are about 540 million years old, while the earth itself, as we've seen, is about 4.55 billion years old. In other words, the fossil-bearing sedimentary beds make up just the last 12% of the Earth's history, and the Precambrian rocks made up the other 88% of the Earth's history. With the ability to determine ages for the Precambrian rocks, geologist finally started making sense of the Precambrian rock record.

Geologic time

Geologists now divide the geologic history of the Earth up into four eons. These are the Hadean (> 4 billion years ago), the Archean (4 to 2.5 billion years ago), the Proterozoic (2.5 billion to 540 million years ago), and the Phanerozoic (540 million years ago to the present.)  Eons are divided into eras, which are further divided into periods, which are divided into epochs. The following table summarizes these divisions of time. You may find it useful to bookmark this table for easy reference as you read the rest of this book. Time before the present in this table is given in units of ka, thousands of years, and Ma, millions of years


Phanerozoic (540 Ma to present)

Cenozoic (66 Ma to present)
The Age of Mammals

Quaternary (2.58 Ma to present)
The Age of Man

Holocene (11.7 ka to present)
Historical Man

Pleistocene (2.58 Ma to 11.7 ka)         

Neogene (23 to 2.58 Ma)

Pliocene (5.3 to 2.58 Ma)

Miocene (23 to 5.3 Ma)

Paleogene (66 to 23 Ma)

Oligocene (34 to 23 Ma)

Eocene (56 to 34 Ma)

Paleocene (66 to 56 Ma)

Mesozoic (252 to 66 Ma)
The Age of the Dinosaurs

Cretaceous (145 to 66 Ma)

Jurassic (201 to 145 Ma)

Triassic (252 to 201 Ma)

Paleozoic (540 to 252 Ma)

Permian (299 to 252 Ma)

Pennsylvanian (323 to 299 Ma)

Mississippian (359 to 323 Ma)

Devonian (419 to 359 Ma)

Silurian (443 to 419 Ma)

Ordovician (485 to 443 Ma)

Cambrian (540 to 485 Ma)

Proterozoic (2500 to 540 Ma)

Neoproterozoic (1000 to 560 Ma)

Mesoproterozoic (1600 to 1000 Ma)

Paleoproterozoic (2500 to 1600 Ma)

Archean (4000 to 2500 Ma)

Hadean (4550 to 4000 Ma)

I have omitted epochs before the Cenozoic Era, periods before the Phanerozoic Eon, and eras before the Proterozoic Eon, since these will not be referenced in this book.

The geologic birth of New Mexico

1.8 billion years ago, during the late Paleoproterozoic, a vast and barren continent lay beneath a leaden sky. This was Laurentia, which would someday become the core of North America. The sun shone low, for the west coast of Laurentia, which would someday become southern Wyoming, lay north of 60 degrees latitude. Compared with its orientation today, Paleoproterozoic Laurasia was rotated clockwise by more than ninety degrees, so that the coastline facing west would face south today. Most of the rest of the Earth's continental crust lay to the south and east, where it had assembled into the supercontinent of Columbia. A broad stretch of ocean, studded with islands, lay to the west.

The surface of the continent appeared devoid of life, for the first true land plants would not appear for another 1.3 billion years. However, life was already present in the coastal waters, and had been since early in the Archean.

Archean life consisted only of bacteria and archaea, two of the three domains of life currently found on Earth. Both have very simple cells that lack true nuclei. Oxygen was virtually absent from the Archean atmosphere. The Sun, having left behind its exuberant infancy, shone at only about 75% of its current brightness, but abundant atmospheric methane trapped enough heat to permit the oceans to remain unfrozen.

The earliest record of life on Earth may be deposits of graphite, the soft crystalline form of carbon used in pencils, in the Isua region of Greenland. These are around 3.8 billion years old. However, this remains a matter of debate, since abiogenic graphite can form from ferrous carbonate at high temperatures. The evidence that at least some of the graphite at Isua is derived from ancient life includes the carbon isotope ratio, 13C/12C, in the graphite. The enzymes of living organisms are selective enough that they react significantly more slowly with molecules containing the 13C isotope than the much more common 12C isotope, and, as a result, biogenic carbon is depleted in 13C compared with the cosmic ratio. Some of the graphite of the Isua region shows such depletion, and, under the electron microscope, the graphite particles are seen to take the form of tubes and granules rather than the flaky grains typical of abiogenic graphite.

Stromatilite fossil
Fossilized Precambrian stromatolites from Glacier National Park. National Park Service

The earliest widely accepted evidence of life dates to some 300 million years later and takes the form of stromatolites. These are distinctively layered rock mounds, around a meter (3') in size, produced by massive colonies of cyanobacteria. Cyanobacteria, formerly known as blue-green algae, are capable of producing oxygen by photosynthesis. Stromatolites have been found in Archean rocks that are 3.5 billion years old. Stromatolites became widespread during the Proterozoic Eon, then declined sharply, likely because other forms of life evolved that feed on the microorganisms making up the colonies. Today, living stromatolites are found only in unusually harsh marine environments in which predators cannot survive.

The oxygen generated by Archean cyanobacteria was removed from the environment as fast as it was generated. It rapidly combined chemically with reduced iron and sulfur dissolved in the ocean water. We know that oxygen levels were very low during the Archean, because Archean river beds preserved in the rock record sometimes contain grains of pyrite or uraninite. Pyrite, FeS2, "fool's gold", is a compound of ferrous iron with reduced sulfur, and while it is common in volcanic rock and in sedimentary rock deposited in low-oxygen conditions, it is unknown in modern river deposits. It quickly oxidizes in today's high-oxygen atmosphere. Uraninite, UO2, likewise rapidly oxidizes to U3O8 under modern conditions.

However, by the start of the Proterozoic Eon, life was beginning to change the face of the earth, as oxygen produced by cyanobacteria exhausted the supply of reduced iron and sulfur and began to accumulate. The accumulating oxygen produced two distinctive geologic signatures, both showing that the ferrous iron (Fe+2) of the young Earth was being oxidized to ferric iron (Fe+3).

Banded iron formation
Banded iron formation from Michigan. Wikimedia Commons

One signature is banded iron formations. These are massive beds of chert (fine-grained silica), magnetite (Fe3O4), and hematite (Fe2O3) in thin layers. Unlike ferrous iron, which is moderately soluble in water, ferric iron is highly insoluble, and it precipitated out of the oceans in large quantities to form the banded iron formations. Banded iron formations are almost always Paleoproterozoic in age, between 2.4 and 1.8 billion years old, and they are now a major source of iron ore. The oldest rocks found in New Mexico are Proterozoic rocks about 1.77 billion years old, and banded iron formation is found in New Mexico in the Tusas Mountains.

The other signature of free oxygen is the presence of sedimentary red beds, which derive their bright red color from hematite. Evidence of Archean red beds has been found in the Timiskaming district of eastern Ontario, but the first widespread occurrence of red beds was around 1.8 billion years ago.

The Proterozoic also marked the emergence of eukaryotes. Bacteria and archea, collectively known as prokaryotes, have cells with a single compartment and little internal structure (other than chlorophyll-bearing membranes in cyanobacteria.) Their DNA, the giant molecule that carries the blueprints of life, is a single circular pair of matching strands that floats more or less freely in the cell compartment.  The circular form is an elegant solution to the problem of how to begin or end the DNA strand: The closed circle has no beginning or end! However, this naked strand of DNA is vulnerable to damage, and its transcription (the process of using the blueprints to manufacture proteins) is slightly unreliable. Transcription in prokaryotes is also subject only to simple controls. This means that prokaryotes mutate rapidly, which is not necessarily a disadvantage for single-celled organisms seeking to occupy as many ecological niches as possible. But it also means that prokaryotes cannot readily specialize, a prerequisite for multicellular organisms. While some prokaryotes form large colonies, these are not characterized by much specialization of roles.

The eukaryotes, by contrast, have cells with a complex internal structure, including multiple separate compartments such as the nucleus. Their DNA is organized into linear pairs of matching strands called chromosomes. These begin and end with telosomes, sections of DNA that bind particularly tightly between pairs to prevent the DNA from unravelling. (Such unravelling seems to be part of the natural aging process of multicellular organisms like ourselves.) The organization of DNA into multiple chromosomes within a special compartment makes for a much lower mutation rate and more sophisticated control of transcription. This in turn makes true multicellular life possible. Thus, modern eukaryotes include all animals, plants, and fungi, as well as several kingdoms of single-celled organisms such as amoebas and ciliates.

Just how eukaryotes emerged is still being worked out by scientists. However, some of the internal compartments of eukaryotic cells (including mitochondria, which power the cell, and chloroplasts, which carry out photosynthesis) are about the same size and shape as bacteria, and they even have their own circular DNA. This strongly suggests that eukaryotes began through endosymbiosis, where small prokaryotes took up residence within a larger prokaryote. The relationship proved beneficial to both, and the pairing evolved into a single organism.  The mitochondria and chloroplasts of modern eukaryotes are now incapable of independent life, since many of the genes encoding their essential proteins have moved to the nucleus of their host cell.

Although the oldest unambiguously eukaryotic fossils, of red algae, are only about 1.2 billion years old, acritarchs first appeared 1.6 billion years ago. Acritarchs were single-celled organisms that were the same size as modern eukaryotic cells, and there are hints of membrane-bound nuclei in some acritarch fossils.  However, acritarchs became extinct around 500 million years ago and their true nature is uncertain. There is evidence that eukaryotes may have emerged even earlier: The earliest traces of organic compounds characteristic of eukaryotic life are found in rocks dating back to around the beginning of the Proterozoic, 2.5 billion years ago. Such traces of organic compounds are known as molecular fossils. However, since almost all eukaryotes contain mitochondria, and the function of mitochondria is to produce chemical energy by oxidation, it seems likely that true eukaryotes emerged in an environment where oxygen was already present. This may well have been in proximity to cyanobacteria colonies that produced a continual supply of the reactive element.

Paleoproterozoic Laurentia was a collage of crustal fragments, most of which had formed during the Archean.  There is much that is still not known about the Archean Eon and about the process of crust formation, but it is widely believed that the continents started as small bits of continental crust. These gradually assembled, sticking to each other when they were pushed together by the oceanic conveyor. It is not known how long this process took, and there are differences of opinion on whether large continents existed yet during the Archean. However, some of the continental crust that formed the heart of Laurentia included possible miogeoclinal sedimentary beds, formed on the passive margin of a continent following rifting. This suggests there were earlier large continents that broke apart.

There is a fair amount of agreement among geologists that, from the start of the Proterozoic on, the continents have assembled into a supercontinent (containing 75% or more of the Earth's continental crust) about every 750 million years or so. The supercontinent then breaks up again. Perhaps this occurs because of periodic changes in the pattern of convective flow in the mantle that shifts the locations of the mid-ocean ridges. The supercontinent itself may help trigger such changes, by trapping heat in the underlying mantle. While geologists disagree over whether a supercontinent existed during the Archean, there is good evidence that a supercontinent assembled around 1.8 billion years ago. This supercontinent has been given many names: Nuna, Hudson, Protopangea, Columbia.

Basement rocks of North America
Precambrian provinces of North America. U.S. Geological Survey.

Archean rocks form the cores of the modern continents, which are known as the continental shields. Proterozoic rocks underlie much of the sedimentary rock in the continental platforms that surround the shields. Together the platforms and shields form the stable continental cratons.

No Archean rocks are found in New Mexico, because New Mexico didn't yet exist.

North America seems to have begun assembling out of smaller fragments of crust, which geologists call provinces, about 2.0  to 1.8 billion years ago. The largest of these fragments was the Superior Province, which took in the Great Lakes area and the adjoining parts of central and eastern Canada. Another was the Slave Province of northwest Canada, which had previously assembled out of three smaller provinces. At around 1.84 billion years ago, these provinces collided and merged to form  Laurentia. Shortly afterwards, two more provinces merged with Laurentia to form the future Wyoming and Montana region.  By 1.8 billion years ago, the western margin of Laurentia, which would face south today, ran roughly along what is now the Wyoming-Colorado border.

The difference between compass orientation today and in the past is bound to complicate our story. Rather than constantly saying things like "west, which would be south today", I'll tell the story as if New Mexico was in its modern orientation, with only occasional comments on what the actual orientation was at the time.

Precambrian provinces of
        western U.S.
Precambrian provinces of western North America. U.S. Geological Survey. Click to enlarge

During the next three hundred million years or so — longer than the time interval between the emergence of the first reptiles and the present day — a mid-ocean ridge was active south of Laurentia. The oceanic lithosphere spreading from this ridge subducted under the southern margin of Laurentia, and a sequence of microcontinents and oceanic island arcs carried by the oceanic crust were brought up against the continent.

A microcontinent is a small patch of continental lithosphere. The largest modern examples are Madagascar and New Zealand, but microcontinents can be as small as individual islands like Socotra. It is possible that most of the continental crust of the Earth started out as microcontinents, which assembled to form large continents.

As we saw in the last chapter, oceanic island arcs are formed when oceanic lithosphere subducts under oceanic lithosphere, as is the case with many of the island chains of the western Pacific. That this is most common in the western Pacific, where the ocean basin is far from the mid-ocean ridge, suggests that this is a phenomenon of cold oceanic crust. This crust more easily subducts. The island arcs above these subduction zones consist of rock that is less dense than oceanic crust, but more dense than typical continental crust.

When microcontinents or island arcs are carried into a destructive margin by the motion of the underlying oceanic lithosphere, they are unable to subduct because of their low density. If the microcontinent is quite small, the lighter crust shears off the underlying upper mantle and sticks to the continent on the other side of the subduction zone. Accumulated sheared crust forms an accretionary wedge on the continental margin. The confused mass of broken and deformed rock beds in the wedge are spoken of as a melange. The Franciscan Complex of coastal California is an example of a melange.

When a large microcontinent is drawn into a subduction zone, the collision is more violent, throwing up high mountains on both sides of the collision zone, which geologists call a suture. We see this process taking place in the Himalayas today. India was once a large microcontinent (or small continent, depending on where you choose to draw the line) that was carried into the southern coast of Asia and is now sutured to the Asian continent along the line of the Himalayas. The collision event itself is known as an orogeny, from the Greek ὄρος oros, "mountain" + γένεσις genesis for "creation, origin". The zone of deformed crust and mountain building along the suture is called an orogen

The microcontinents and island arcs that merged with southern Laurentia between 1.7 and 1.6 billion years ago formed the Yavapai and Mazatzal Provinces, which reach from modern southwest Arizona to Michigan and includes most of Colorado and New Mexico. The large region of accreted material making up these provinces consists of a mosaic of blocks between about 10km and 100 km (6 to 60 miles) in size, separated by shear zones, which are narrow zones of high deformation. The oldest Precambrian rocks in northern New Mexico belong to Yavapai formations that are about 1.77 billion years old.

Because they represent two episodes of a long process of accretion along the southern boundary of Laurentia, the Yavapai and Mazatzal Provinces are sometimes grouped together as the Transcontinental Proterozoic Provinces. This accretion process is one of the most significant crust forming events discernible in the geologic record. Some geologists have suggested that the best modern counterpart is southeast Asia, where the island arcs of Indonesia and Malaysia are being welded onto Asia as a result of subduction both on the south (Indian Ocean) and east (Pacific Ocean).

The Yavapai and Mazatzal Provinces and the Jemez Lineament

If you examine a geologic map of the southwest United States, you will find a line of young volcanic fields stretched across New Mexico and Arizona. These include the Raton volcanic field, the Mora volcanic field, the Taos plain, the Jemez Mountains, Mount Taylor, the Lucero volcanic field, the Zuni-Bandera volcanic field, the Springerville volcanic field, the White Mountains volcanic field, and the San Carlos volcanic field.

When plate tectonics was still quite new, geologists identified the Snake River Plain as a hot spot trace.  Volcanoes repeatedly erupted over a fixed point in the deep mantle as the North American plate moved southwest over this mantle hot spot. This made a great deal of sense, since the youngest volcanoes are at Yellowstone (and are potentially still active) and the oldest, most thoroughly extinct volcanoes are far to the southwest, in northern Nevada. Geologist today still believe the Snake River is a hot spot trace, though there is debate among geologists about the exact nature of the hot spot.

The Jemez Lineament seemed to fit the same pattern. The volcanism followed a path of similar length and direction, and some of the volcanoes at the northeast end of the Lineament were obviously very young. However, as the rocks along the lineament were radiometrically dated, the hot spot theory for the Jemez Lineament began to fall apart. There is no systematic progression in age along the lineament. Volcanism began a little earlier towards the center of the lineament, but quickly spread southwest and northeast. This is not consistent with a hot spot.

It is now widely believed that the Jemez Lineament is an ancient structure of some kind in the lower crust or upper mantle. The most widely accepted explanation is that the Jemez Lineament marks a hydrated subduction zone scar where, some 1.7 billion years ago, active subduction took place along what was then the continental margin. Accretion of additional island arcs then jammed the subduction zone and active subduction shifted further south.The relic subduction zone contains minerals with an unusually low melting point, making it a fertile source rock for production of magma.

The Jemez Mountains are located squarely on the intersection of the Jemez Lineament with the western margin of the Rio Grande Rift. The Rift is a region of the crust stretching from central Colorado into northern Mexico, roughly along the valley of the Rio Grande, where the crust began to be slowly pulled apart about 30 million years ago. We'll have much more to say about the Rio Grande Rift later in the book. Deep faults mark the east and west boundaries of the Rift, separating it from adjoining mountain ranges, such as the Tusas, the Sangre de Cristo, the Sierra Nacimiento, and the Sandia Mountains.

The pattern of volcanism is not the only evidence for the existence of the Jemez Lineament. Precambrian rocks north of the lineament have maximum ages of around 1.77 billion years, as we've seen. South of the Lineament, the maximum ages are around 1.7 billion years, and there are subtle differences in isotope ratios. These are interpreted as different isotope model ages for the two regions of crust. The steady decay of 147Sm to 143Nd in the earth's mantle, mentioned in the last chapter, means that fresh magma extracted from the upper mantle is increasingly enriched in 143Nd with the passage of geologic time. The isotope model ages for rocks north and south of the Jemez Lineament show that those south were formed from material extracted from the mantle significantly later.

Another piece of evidence is the strength of magnetic fields measured by aeromagnetic surveys. Sedimentary rocks normally are very poor in magnetic iron minerals, making them "transparent" to magnetic fields. Thus an aircraft flying low over the ground can measure the magnetic field of the basement rock underlying the surface sedimentary beds, except in the immediate area of recent volcanism. Such surveys were originally carried out to find hidden ore bodies, but they also allow geologists to trace basement structures. Magnetic fields are anomalously high along the Lineament, suggesting a buried structure rich in magnetic iron minerals.

Finally, there is seismic profiling evidence for a deep structure coinciding with the Jemez Lineament. Seismic profiling was originally developed by the petroleum industry and is a way to get information about the subsurface rocks using sound waves. It is similar to probing the structure of the Earth using P-waves from earthquakes, but you don't have to wait for a nearby earthquake. Seismic profiling is carried out by setting out a network of seismic detectors and then generating sound waves using explosives lowered into a borehole, by dropping an extremely heavy weight into a borehole, or by placing a large, heavy metal plate on the ground and vibrating the plate at a carefully chosen frequency. Each approach has its advantages. The sound waves are reflected when they hit a boundary between rocks of different types, and careful measurement of the return times of the sound waves can be used to map out the rock layers below the surface. It's like sonar, but for use underground rather than in the ocean.

Seismic profiling of the Jemez Lineament reveals that deep rock beds on either side dip into the Lineament. This is more pronounced on the north side, and the general structure suggests that the Lineament is where the Yavapai plate to the north was overridden by the Mazatzal plate to the south. Some subduction took place, and the hydrated subduction zone scar may well be an excellent source rock for production of magma, because it likely contains abundant hydrous minerals.

The Jemez Lineament appears to have been severely deformed by motion along deep north-trending faults across north-central New Mexico. Most of this deformation likely took place during the Laramide Orogeny, a period of mountain building that peaked around 50 million years ago. Some reconstructions suggest that the Precambrian rocks of the Sierra Nacimiento were originally directly west of the Precambrian rocks of Sandia Crest, while the Precambrian core of the Sangre de Cristo lay well to the north.

The Tusas Mountains and the Spring Creek Shear Zone

... these waters have been recommended by Doctor Nagle, of Santa Fe, in many chronic diseases, and always with success.

— Lieutenant William G. Peck, 1847

North of Espanola, beyond the confluence of the Rio Chama and Rio Grande Rivers, lies Black Mesa. The road to Chama skirts the west end of the mesa, and a side road, U.S. 285, turns along the north side of the mesa and follows the Rio Ojo Caliente to the village of the same name. Ojo Caliente, "Hot Pool" in Spanish, is the location of hot springs near the river. The Spanish discovered the springs early in their settlement of New Mexico, but the exposed location, subject to Comanche raids, prevented permanent settlement until 1868. In that year, Antonio Joseph, the first Territorial representative to the U.S. Congress, built a bathhouse at the springs. This has grown into a small resort favored by the gentry of Santa Fe.

West of Ojo Caliente is a ridge of ancient rock, and to the northwest is Cerro Colorado, "Red Hill". These are the southernmost outliers of the Tusas Mountains, which separate the San Luis Valley and Taos area to the east from the Chama Valley to the west. Cerro Colorado is covered with pinon scrub forest, like much of the surrounding area, but as one proceeds north, the scrub gives way to ponderosa pine forest. Separate roads lead to La Madera and Tres Piedras, the gateways to the Tusas.

Ojo Caliente is located just beyond the northern boundary of the Jemez region as depicted in our digital maps, and the northern Tusas Mountains well beyond that. However, the rocks of this region tell a story that is crucial to our understanding of the Jemez Mountains, so we will briefly venture north.

The Tusas Mountains are underlain by Precambrian rocks of the Yavapai and Mazatzal Provinces. The precise boundary between the two has proven difficult to pin down, but there is a zone 300 km (200 miles) wide that seems to be transitional between the two provinces. The southern edge of the transitional zone is roughly coincidental with the southern edge of the Jemez Lineament. Its northern edge is unusually sharp and well exposed in the Tusas Mountains and has been thoroughly studied by geologists interested in the process of continent assembly. This boundary is defined by a lithological discontinuity where rocks assigned to the Yavapai Province north of the discontinuity abruptly give way to rocks assigned to the Yavapai-Mazatzal transition zone south of the boundary. The lithological discontinuity is nearly coincident with a major structural feature, the Spring Creek Shear Zone, which, unsurprisingly, lies along Spring Creek. This feature is probably younger than the rock beds themselves, having likely formed around 1.4 billion years ago.

The Precambrian rocks of the Tusas Mountains have been distorted and altered by geologic processes over the last 1.77 billion years, a process called metamorphosis.

Geologists divide rocks into three large families. Igneous rocks form directly from magma. They include such rock types as granite, which forms from silica-rich magma that hardens underground; basalt, which forms from silica-poor lava that erupts at the surface; and ignimbrite, which forms from hot volcanic ash. Sedimentary rocks form from beds of clay, sand, pebbles, or other fragments of older rock, or of minerals precipitated from large bodies of water, that are gradually cemented together, usually by additional minerals precipitated from ground water. They include rocks like sandstone, shale, and limestone. Metamorphic rocks form from existing rock when it is subject to heating that causes the rock to recrystallize without actually melting.

The heat and pressure required to form metamorphic rock is usually found only deep underground. When metamorphic rocks are found near the earth's surface, they are a strong indication that tectonic forces have brought up rock that was once deeply buried, a process geologists call exhumation. No, really.

Exhumation is possible because of isostasy. Isostasy is the term for the balance between the weight of the mountains and the buoyancy of the thickened crust beneath them. As the mountains are worn down by erosion, uplift raises new mountains to restore the balance. When you consider that continental crust underneath the Himalayas is around 100 km (60 miles) thick, versus 40 km (25 miles) for more normal continental crust, it is not hard to see that prolonged erosion of a high mountain range can bring rock to the surface that was originally very deep underground.

Metamorphic rocks are sometimes classified by the original igneous or sedimentary rock from which they formed (their protolith). Thus one speaks of metarhyolite, metabasalt, or metaconglomerate if it is possible to determine that the protolith was rhyolite, basalt, or conglomerate. However, as the degree of metamorphism increases, the original form of the rock becomes hard to discern. Such rocks are classified according to their mineral content and degree of foliation. A foliated metamorphic rock is one in which the minerals have segregated into distinct bands. Foliation shows the direction in which stresses were applied to the rock while it was undergoing metamorphosis, with the foliation typically lying perpendicular to the direction of greatest compression.

The mineral content of a metamorphic rock gives clues to the temperature and pressure at which the rock underwent metamorphosis. This is because different minerals are stable under different conditions. Geologists speak of characteristic combinations of minerals that point to particular temperature and pressure regimes as metamorphic facies. I won't go into these in any detail, because metamorphic rocks are uncommon in the Jemez. The Jemez is mostly composed of relatively young rocks that have not experienced deep burial.

The Moppin Complex

North of Spring Creek is Hopewell Ridge, which is composed mostly of rocks assigned to the Moppin Complex. These are among the oldest rocks found in northern New Mexico, with a radiometric age in excess of 1.75 billion years, and are typical of the Yavapai Province. Similar rocks are found around Gold Hill in the Sangre de Cristo Mountains and are thought to be part of the same ancient block of crust, but displaced by motion along the Rio Grande Rift. The Moppin Complex consists of thick beds of mafic volcanic rock interbedded with occasional thinner beds of felsic volcanic rock and sediment. All have been metamorphosed.

A particularly interesting feature of Hopewell Ridge is the presence of magnetite schist. This was once prospected as iron ore, but mining a limited quantity of ore so far from existing rail lines is not economical.

Banded iron
Magnetite schist.  36 38.331N 106 8.072W

Magnetite schist is probably metamorphosed banded iron formation. The presence of banded iron formation on Hopewell Ridge is one indication that the Moppin Complex rocks were erupted in a marine environment, as part of an island arc. Another indication is the presence of well-preserved pillow basalt, erupted under water, in Moppin Complex outcrops in the Brazos Peak area. The isotope model age of the Moppin Complex is close to its crystallization age, indicating that the rock is juvenile, formed from magma freshly extracted from the mantle rather than from reworked older crust. This distinguishes Yavapai rock from the Archean core of Laurentia, where isotope model ages are often much greater than crystallization ages, and is another indication that the Yavapai Province formed by accretion of island arcs.

The Moppin Complex is bimodal, meaning that the volcanic rocks and sediments from which it formed included high silica and low silica magmas, but little intermediate magma. An exposure of feldspathic schist along Hopewell Ridge is an example of a felsic member of the Moppin Complex.

Granite gneiss

Granite gneiss
Feldspathic schist. Near 36 38.066N 106 7.646W

A feldspathic rock is composed mostly of fine grains of quartz and feldspar. These are visible under the loupe, which also shows smaller quantities of a mafic mineral, possibly mica or amphibole. The rock here is schistose, having a laminated structure as shown by the thin layers of the mafic mineral, and hence is described as feldspathic schist.

The minerals in feldspathic schist are important characters in our story, deserving of proper introductions.


Quartz is a mineral composed of silicon dioxide, SiO2. We've seen quite a bit of quartz already, but we'll now examine this important mineral more closely.

Silicon atoms prefer to covalently bond with four oxygen atoms. Each of these oxygen atoms shares a pair of electrons with the silicon atom, allowing the silicon atom to surround itself with a shell of eight electrons. This is a particularly stable structure for most light chemical elements. Each oxygen, in turn, prefers to covalently bond to two silicon atoms, which likewise allows the oxygen atom to surround itself with a shell of eight electrons. (Two pairs of electrons are shared with silicon atoms, and two the oxygen keeps to itself.) If we were living in a Flatland world of two dimensions, a quartz crystal might form as shown in the following diagram:

Electron dot diagram of silica crystallization
Electron-dot diagram of the formation of a hypothetical 2-D silica crystal

Each isolated silicon atom starts out with four outer shell electrons, and each isolated oxygen atom starts out with six outer shell electrons. When these atoms bond together to form quartz, the atoms in the interior of the quartz crystal all end up surrounded by the ideal shell of eight electrons.

Of course, we don't live in a two-dimensional world, and a real quartz crystal has a much more complicated three-dimensional structure. The four oxygen atoms bonded to each silicon atom lie at the corners of a tetrahedron, not in a flat plane. Nor do the two silicon atoms bonded to each oxygen atom form a straight line. Instead, because each pair of electrons in a filled electron shell wants to lie at a corner of a tetrahedron, the two pairs shared by silicon atoms lie at an angle close to 144 degrees rather than 180 degrees. (The angle is not the ideal 110 degrees of a tetrahedron, because the two silicon atoms repel each other enough to distort the tetrahedron.) This means that two silica tetrahedra sharing an oxygen atom lie at an angle of 144 degrees to each other. The tendency of the silica tetrahedra in quartz to find an arrangement in which the tetrahedra all lie at 144 degrees to each other is part of the reason for the peculiar structure of a quartz crystal, which is quite hard to visualize from two-dimensional images. 

Nevertheless, I'll make an attempt here to explain the quartz structure, since quartz is so important. We'll start by examining the unit cell, which is the smallest piece of any crystal that contains the basis of its entire structure. A unit cell is always a parallelepiped; that is, it is a volume of space bounded by six faces with opposite faces parallel. For example, a cube is a parallelepiped in which the sides are squares meeting at right angles.  In quartz, the unit cell has top and bottom that meet the sides at right angles, but the sides meet at angles of 60 and 120 degrees. The entire structure of a crystal can be generated from its unit cell simply by packing copies of the unit cell together so the faces all line up.

The unit cell of quartz is deceptively simple.

Unit cell of quartz
Unit cell of alpha quartz

Each silicon atom is represented by a gray sphere and each oxygen atom by a red sphere, with the bonds shown as sticks joining the spheres. The spheres are not to scale, being shrunk down in size to show the bonds better; the spheres would be in contact in a scaled depiction. There are thee silicon atoms and six oxygen atoms in the unit cell.

I know: You see six silicon atoms in the diagram. But the silicon atoms all lie on the faces of the cell, and so are shared with the neighboring cells. We could shift the boundaries of the unit cell so that our diagram shows just three silicon atoms -- the unit cell definition is not unique -- but this would not display the structure as well. The diagram shows bonds extending from the silicon atoms on the cell faces into the neighboring cells. If you examine the diagram for a few moments, you should be able to convince yourself that the pattern does indeed repeat itself, with (for example) the silicon atom on the top matching the silicon atom on the bottom. The silicon atom on each face has two bonds extending into the unit cell and two bonds extending into a neighboring cell. This makes the structure equally strong in all directions.

It can be startling to discover how this simple unit cell generates a wonderfully complicated crystal structure. To illustrate, we're going to show a single layer of unit cells, generated by lining up unit cells side to side and leaving the top and bottom free. Looking down on this layer, we see:

Single layer of
A single layer of alpha quartz

The unit cells are marked in this image. The full crystal consists of stacks of layers identical to this one. The silicon atom appearing as a small black sphere at the center of each unit cell is actually two silicon atoms, one on the top and one of the bottom face, that are vertically superimposed. These link the layers in the crystal.

The diagram shows that there are large channels running the length of the crystal; one such channel is marked in the version below.

Single layer
          of quartz emphasizng channel
A single layer of alpha quartz with one of the channels outlined

Because of these channels, a quartz crystal has a fairly open structure. This gives quartz a relatively low density, about 2.65 grams per cubic centimeter. (For comparison, the density of water is almost exactly 1.0 grams per cubic centimeter.) However, the strong three-dimensional bonding gives quartz the greatest hardness of any common mineral. Quartz is also chemically inert and very stable under the conditions found at the surface of the earth.

The Internet Quartz Page has additional information on the wonderful and complicated structure of quartz.


Feldspar is a mineral that is similar in structure to quartz, but some of the silicon atoms have been replaced with aluminum atoms. An aluminum atom has one less electron than a silicon atom, and the missing electron must somehow be supplied if an aluminum atom is to take the place of a silicon atom in the crystal structure. Returning again to our Flatland world, the formation of a feldspar crystal might take place as:

Electron dot diagram of silica crystallization
Electron-dot diagram of the formation of a hypothetical 2-D microcline crystal

A silicon atom has been replaced with aluminum, and a nearby potassium atom provides the missing electron needed to complete the structure. The potassium atom fits snugly into one of the openings in the structure, near the aluminum atom to which it donated its electron. As with quartz, the structure of a real feldspar in our three-dimensional world is much more complex and quite difficult to visualize from two-dimensional images. It is also not simply the quartz structure with added potassium; the silica and alumina tetrahedra still form a three-dimensional structure, but one that is subtly different from quartz, giving the potassium a little more room to fit in the structure.

Atoms of sodium also readily donate an electron, while a calcium atom can provide two extra electrons to two aluminum tetrahedra. This gives us the three most common varieties of feldspar: potassium feldspar, KAlSi3O8; sodium feldspar, NaAlSi3O8; and calcium feldspar, CaAl2Si2O8.

I have described the bond between oxygen and silicon as covalent, because the bond consists of a pair of electrons shared between the two atoms. However, this is an idealization, like many things in science. Oxygen is tremendously greedy for electrons: Only the much less common element, fluorine, has a greater electron affinity. So the sharing is unequal, with the electrons being more tightly bound to the oxygen than the silicon. With aluminum, the sharing is even more unequal. With other metallic elements, the sharing is so unequal that the electrons effectively have been lost to the metal and belong to the oxygen atom. Such a bond is called an ionic bond, because both atoms have been ionized: The metal atom, shorn of one or more of its electrons, now has a net positive charge (making it a cation) while the oxygen atom, having acquired two electrons from its neighbors, has a net negative charge (an anion.) The ions are bound to each other because of their overall opposite charges. Geochemists find it convenient to speak of all atoms in a crystal as if they have been ionized, even when the bonding has considerable covalent character, as with oxygen and silicon. I will follow this convention from here on.

Potassium feldspar comes in three separate varieties, or polymorphs, each of which is stable in a different range of temperature and pressure. The form stable at low temperature is called microcline. 

          feldspar from the Harding mine
MIcrocline feldspar from the Harding Mine. Feldspar of this quality is rare in the Jemez. 36 11.557N 105 47.695W

Orthoclase is stable at elevated temperature, and sanidine becomes the stable form at the highest temperatures. The high temperature polymorphs are not uncommon in nature, because rapid cooling after their formation can freeze the crystal structure before it has time to convert to a lower temperature form. The conversion from one polymorph to another can be thought of as a kind of chemical reaction, and like many chemical reactions, it takes place only at high temperature.

Potassium feldspar is often found in the same rocks as quartz, but it is easily distinguished by its tendency to fracture along flat surfaces at nearly right angles, as in the photograph above. This property is called cleavage. The number and relative angles of cleavage planes are characteristic of any mineral. Quartz has no cleavage planes, breaking instead along irregular curved surfaces like those of thick broken glass. In addition, quartz is usually nearly colorless and transparent while potassium feldspar is translucent and often has a pink to brick red color.

Calcium and sodium freely substitute for each other in feldspar, forming what geologists call a solid solution series. This is because of the similarity in the sizes of sodium and calcium ions. The sodium ion has a radius of about 0.97 Angstroms (0.97 x 10-8 meters). The calcium has a very similar radius of 0.99 Angstroms. This is about 70% of the radius of an oxygen ion. Both ions fit very nicely into a site in the feldspar structure that is surrounded by eight oxygen ions. Because it has almost the same radius, a calcium ion easily substitutes for a sodium ion, so long as an aluminum ion simultaneously substitutes for a silicon ion to maintain charge balance.  Calcium-sodium feldspar is called plagioclase, and plagioclase with all compositions from nearly pure sodium feldspar (albite) to nearly pure calcium feldspar (anorthite) is found in nature. Plagioclase can often be distinguished from potassium feldspar because its cleavage surfaces are striated, or marked by very fine parallel grooves.

Potassium does not easily substitute for calcium or sodium, because its ions are significantly larger, at 1.33 Angstroms. It can just fit into the feldspar structure, if the structure is distorted to make more room for the potassium ions. In sanidine, sodium substitutes fairly freely for potassium, but if the feldspar cools slowly enough to convert to orthoclase, the sodium tends to separate out into thin layers of albite to give what is called perthitic feldspar. Most microcline is perthitic.

Ion size also explains why there is no such thing as magnesium or iron feldspar. Both metals readily donate two electrons, like calcium, and it seems like they might be able to replace calcium in feldspar. However, the magnesium ion (with a radius of 0.66 Angstroms) and ferrous iron ion (with a radius of 0.64 Angstroms) are significantly smaller than potassium, calcium, or sodium ions. Ferrous iron and magnesium prefer to be surrounded by just six oxygen ions, which is not possible in the feldspar structure. However, small amounts of ferric iron (radius 0.63 Angstroms) can substitute for aluminum (radius 0.53 Angstroms) in potassium feldspar, with some distortion of the structure. This trace of iron gives most potassium feldspar its characteristic pink to brick red color.

The remaining components of our feldspathic schist outcrop are mafic minerals, mica and amphibole. Mafic minerals are minerals rich in iron and magnesium, and they tend to be dark in color.

A composition of quartz and feldspar with smaller amounts of mafic minerals is characteristic of granite, of which we'll see some beautiful examples later in this chapter. The feldspathic schist shown earlier has this granite-like composition, and these minerals are characteristically separated into layers in the rock. This thin layering suggests the presence of muscovite mica, which in turn is an indication of abundant aluminum in the rock. This suggests either an aluminum-rich granite protolith or a sedimentary protolith rich in clay, such as shale, of which we'll see many examples later. The thin layering is typical of shale and may indicate that this is actually a metashale.

The Maquinita Granodiorite and calk-alkaline and other igneous suites

Another clue to the history of northern New Mexico in the Precambrian is the presence of calc-alkaline igneous rocks within the Moppin Complex. The most widespread in the Tusas Mountains is the Maquinita Granodiorite, which has been dated at 1.755 billion years old.


Maquinita Granodiorite

Granodiorite is an intermediate-felsic intrusive igneous rock. An  intermediate-felsic rock is an igneous rock with a fairly high silica content, between 63% and 69%, like dacite. An intrusive igneous rock is a rock that solidifies from magma that is trapped underground. Because the surrounding solid rock is an excellent insulator, the magma cools extremely slowly, and there is time for relatively large crystals to form. These are easily visible with magnification and are often obvious even to the naked eye. In a granodiorite, the crystals are found to be quartz and feldspar with some mafic minerals, much like granite. However, the feldspar in granodiorite is mostly plagioclase rather than alkali feldspar, which is the more abundant feldspar in granite.

The significance of the Maquinita Granodiorite is that, in addition to having a fairly high silica content, it is also moderately enriched in the alkali metals, potassium and sodium, and the alkaline earths, calcium and magnesium. Rocks that are enriched in this way, and which show other distinctive chemical characteristics (such as a high aluminum content and a tendency to steadily decrease in iron content as the silica content increases) are described as calc-alkaline

Geologists speak of igneous suites, which are families of igneous rocks having a similar origin. Each suite comes from its own distinctive source rock subject to a particular degree and type of partial melting. Calc-alkaline magma tends to form from rocks that have already experienced some partial melting (moderately depleted source rocks) in an environment that is more oxidized and contains more water vapor than is the case with the other common suite, the tholeiitic suite. The water vapor alters the eutectic compositions, and the relatively high content of oxygen means that, as the magma differentiates, much of its iron is removed as magnetite crystallizes out. By contrast, tholeiitic magma is poor in oxygen, and as it differentiates, the iron content actually increases as a magnesium-rich mineral called olivine crystallizes out, instead of iron-rich magnetite.

The calc-alkaline family of rocks are characteristically erupted over subduction zones, where fluids "sweated" from the subducted slab provide water and oxygen, and the production of magma from the mantle wedge rapidly depletes the source rock. The presence of calk-alkaline rocks in northern New Mexico Precambrian formations further reinforces the idea that the Yavapai Province formed by accretion of island arcs along a destructive margin.

Both the calc-alkaline and the tholeiitic suites are described as subalkaline. Subalkaline rocks are notable for being silica saturated, meaning that there is enough silica in the rock for its entire alkali metal content to form feldspar. By contrast, alkaline rocks have a high enough content of alkali metals that they are silica undersaturated, so that some of the alkali metals are present as silicate minerals with a lower silica content than feldspar. Alkaline magmas are through to be produced at a greater depth or from a lower degree of partial melting than tholeiitic magmas. We'll have more to say about silica saturation in a later chapter.

The Vadito Group

The Spring Creek Shear Zone neatly divides rocks of the Yavapai Province to the north, which are assigned to the Moppin Complex, from younger rocks of the Yavapai-Mazazatl transition zone to the south, which are assigned to the Vadito and Hondo Groups.

Throughout this book, you'll find rocks identified by their group, formation, or member. For example, the Bandelier Tuff is one of the most important formations in the Jemez area. It names a distinctive kind of volcanic rock found throughout the Jemez that was formed by two similar caldera eruption events 1.25 and 1.61 million years ago. This formation is divided into the Tshirege Member and the Otowi Member, corresponding to the two individual events. The Bandelier Tuff is one of several formations making up the Tewa Group, which includes most of the rock erupted in the Jemez in the last two million years. Much of this book is organized around describing formations in decreasing order of age.

One can subdivide members into beds and combine groups into supergroups. We will mostly refrain from doing so in this book. The important thing to remember is that a group consists of related formations, which in turn consist of related members. When a formation or member is composed almost entirely of a single rock type, it is described using that type, as with the Maquinita Granodiorite or the Bandelier Tuff.

A complex, such as the Moppin Complex, is a body of rock that has been so distorted by metamorphism or igneous intrusion that one can no longer assume that the rock beds are ordered by age, with the younger beds at the top and the older at the bottom. Many of the beds of the Moppin Complex have been so heavily folded that the older beds now lie atop the younger beds.

With that digression on stratigraphy out of the way, let's return to our story.

About 1.71 billion years ago, the island arcs that had formed south of Laurentia began accreting onto the continental margin. This event, known to geologists as the Yavapai Orogeny, would continue for another 30 million years. Accretion was interrupted at least once by the formation of a back-arc basin.

Diagram of subduction zone showing back arc basin

A back-arc basin forms in the crust above a subducting plate. It may be caused by trench rollback, in which the trench marking the point of subduction shifts in the direction of the subducting plate. This stretches the overriding plate, sometimes rifting the plate apart and forming what amounts to a very small ocean basin behind the plate. Back-arc basins tend to close up again, and this process may have taken place, possibly more than once, during subduction under the Yavapai Province. Evidence for the formation of back-arc basins around this time is provided by beds of pyrite-bearing chert near Wheeler Peak in the Sangre de Cristo Mountains. The pyrite and chert are thought to have formed in hydrothermal systems along the axis of the basin.

The basin became a trap for sediments. The first sediments to accumulate were interbedded with some felsic volcanic rocks and contained a fair amount of silt and clay. These beds are assigned to the Vadito Group, which consists mostly of micaceous schist, conglomerate, dirty quartzite, and metarhyolite. All are close to 1.7 billion years in age.

The rock beds of the Vadito and Hondo Group are severely faulted and deformed, to the extent that geologists were for a long time uncertain whether the Hondo Group or the Vadito Group was older. However, it is now reasonable clear that the Vadito Group is slightly older. Among its oldest beds are aluminum-rich schists, which were once mined for kyanite on Mesa de la Jarita.

Kyanite mine on Mes
          de la Jarita
Kyanite mine on Mesa de la Jarita. 36 32.657N,106 04.920W

Kyanite. 36 32.657N,106 04.920W

Kyanite is aluminum silicate, Al2SiO5. The best samples have a striking pale blue color, of which there is just a hint in these samples. Just as potassium feldspar has three polymorphs, so aluminum silicate has three polymorphs; kyanite is stable at lower temperatures and high pressures, which is an indication of the metamorphic conditions where this rock recrystallized. Overlying the aluminum-rich shale beds is a striking conglomerate, the Big Rock Conglomerate.

Big Rock Conglomerate
Big Rock Conglomerate.
36 32.944N 106 05.654W

Notwithstanding its name, this rock is best described as a metaconglomerate. Conglomerate is a sedimentary rock containing a significant quantity of rounded pebbles (clasts) with a diameter of 2mm (0.08 inch) or greater, typically embedded in a sandy matrix. Metamorphosis can transform this rock into metaconglomerate by converting the matrix to quartzite. When this happens, the rock tends to fracture straight through the pebbles, rather than around them as is typically the case in an unaltered conglomerate.

The Big Rock Conglomerate is highly foliated, showing it was strongly compressed and deformed. However, there are numerous large quartz pebbles in the conglomerate which are almost undeformed. This shows the rest of the rock was much softer than the quartz. This likely was gravel in river channels in a muddy floodplain, probably close to volcanoes erupting silica-rich ash.

The Marquenas Quartzite

In the Tusas Mountains, the Big Rock Conglomerate transitions to less spectacular beds of metaconglomerate and micaceous quartzite. An outcropping of metaconglomerate can be found near the forest road on the north rim of Spring Canyon.

An outcropping of Vadito metaconglomerate

Close up of Vadito metaconglomerate
Metaconglomerate bed of the Vadito Group. Kiowa Mountain in the background. 36 36.775N 106 5.532W

To the east of Ojo Caliente are the Picuris Mountains, a part of the Sangre de Cristo Range. Here the Vadito Group is exposed again, and includes some beautiful metaconglomerates.

Metaconglomerate of the Vadito Group. Picuris Mountains.

The photograph is of a large sample now gracing my yard. The original outcropping is quite extensive and is striking, looking like a dry river bed that has been spray painted with gold paint. 

Metaconglomerate of the Vadito Group. Picuris Mountains. 36 12.220N 105 48.424W

The clasts are quite large, well-sorted, and well-rounded. The deposit was subsequently deeply buried and subjected to metamorphosis, to the point that the clasts have been deformed so that they are all flattened in the same direction. The luster is probably from sericite, which is a particular fine-grained form of muscovite mica, KAl2(AlSi3O10)(OH)2. Muscovite is a common mineral in both igneous and metamorphic rocks.


Quartz and feldspar, together with all other silicate minerals built on a basic three-dimensional network of interlocked silica and alumina tetrahedra, are called tektosilicates. Muscovite belong to a different family of silicate minerals, called phyllosilicates. In a phyllosilicate, the silica tetrahedra are joined at only three of their corners, forming sheets of tetrahedra. Unlike the structure of quartz or feldspar, which is tough to depict in a two-dimensional image, it is easy to depict the structure of a phyllosilicate:

Phyllosilicate backbone

Phyllosilicate backbone structure

This graphic is drawn from the perspective of someone looking directly down on a sheet of silica tetrahedra. Three of the oxygen ions in each tetrahedra are shared; the fourth sits by itself at the tip of each tetrahedron, as shown here. This fourth oxygen ion is described as an apical oxygen ion. The overall structure is of layers of interlinked rings of silica tetrahedra.

From a chemical standpoint, this structure is incomplete. The apical oxygen ions are only connected to one silicon ion. In addition, in muscovite, one silica tetrahedron in four is replaced by an aluminum tetrahedron, which makes the structure even more negatively charged. As with feldspar, the negative charge is balanced by metal cations.

Muscovite structure
Muscovite structure. U.S. Geological Survey

Muscovite is composed of layers of triple sheets. The upper and lower sheet of each layer is a phyllosilicate sheet. The sheets are oriented so that they face each other, with the apical oxygen ions on the inside. Between the phyllosilicate sheets is a sheet of aluminum hydroxide, Al(OH)3, a substance which in pure form crystallizes as the mineral gibbsite. A gibbsite sheet is not unlike a phyllosilicate sheet, but with rings of aluminum ions joined by hydroxyl (OH-) ions such that each aluminum ion is bonded to six hydroxyl and each hydroxyl to two aluminum ions. The apical oxygen ions of the outer phyllosilicate sheets bond to the gibbsite sheet by replacing some of the hydroxide ions. It's almost like a sandwich, with the two phyllosilicate layers as the bread and the gibbsite as the sticky layer of peanut butter or marmite that holds the two slices of bread together.

The gibbsite sheet starts out electrically neutral, while the two phyllosilicate sheets start out negatively charged because of their apical oxygens. Each hydroxyl ion that is replaced by an apical oxygen ion reduces the negative charge, and for two pure silica phyllosilicate sheets fully bonding to one gibbsite sheet, it all balances out to make the triple layer neutral. Stacks of such triple layers make up the uncommon mineral, pyrophyllite.

In muscovite, one in four silicon ions in the phyllosilicate sheets are replaced by aluminum, giving the triple layer a net negative charge. Each ring in the phyllosilicate layer forms a kind of cup in the outer surfaces of the triple layer, which is lined with oxygen and hydroxyl ions. This negatively charged cup is an inviting location for a potassium ion to sit. The neighboring triple layers have corresponding cups that fit to the potassium ions and bind the sheets together, producing a structure with no net charge. The binding of triple layers by potassium makes muscovite significantly harder than pyrophillite.

The family of phyllosilicate minerals which share the three-layer structure of muscovite are known as micas. Different mica minerals substitute different metals for aluminum and potassium, but have the same basic structure.

The binding by potassium ions is not particularly strong. As a result, mica is easily split between triple layers. This gives mica crystals a single perfect cleavage plane. It is possible to split mica into very thin sheets, which have been used for insulation, as a dielectric in electronic components, and even as a substitute for glass.

Muscovite from the Joseph Mine. 36 19.646N 106 03.324W

Muscovite is an aluminum-rich mineral, with equal numbers of aluminum and silicon ions in its structure. This contrasts with alkali feldspar, which has three silicon ions for every aluminum ion. The presence of muscovite in granite is a indication that the granite is peraluminous, rich in aluminum. Muscovite in a metamorphic bed suggests the protolith was enriched in clay, which has a high aluminum content.

The mica-rich metaconglomerates of the Marquenas Quartzite in the Picuris Mountains give way to quartzite. Here's a sample.

Ortega Quartzite of
        the Picuris Mountains
Quartzite from the Picuris Mountains, with some magnetite content

The dark layer in this sample contains a small amount of magnetite, which imparts the dark color. When a piece from this layer is crushed, it is found to be mostly colorless quartz grains with a small percentage of much smaller black magnetite grains, which can be separated out with a kitchen magnet.


Iron readily contributes two electrons to the chemical compounds it forms, and iron in this state is known as ferrous iron to geologists. Chemists speak of such iron as having an oxidation number of +2, which is also the charge of a ferrous ion. With a little coaxing, the ferrous ion can contribute a third electron as well, forming ferric iron, with an oxidation number of +3. Both are found in the earth's crust today, but ferrous iron predominated in the early Earth, before cyanobacteria began generating oxygen. Because oxygen is scarce even today in the depths where magma is generated, ferrous iron is significantly more abundant in most igneous rocks than ferric iron.

Magnetite is a bit of a funny critter. Ferrous oxide has the composition FeO, since the two electrons donated by each iron ion match the two electrons needed by each oxygen ion. Ferric oxide has the composition Fe2O3, reflecting the additional oxygen needed to accept the third electron from each iron ion. Magnetite has the composition Fe3O4, suggesting that the iron in magnetite is in a kind of halfway state, with an average oxidation number of +2.5. One can think of magnetite as having the composition (FeO)(Fe2O3). The evidence from crystallography is that there really are separate ferrous and ferric ions in the crystal lattice, each occupying their own pattern of sites, and by a quirk of chemistry, this structure is unusually stable.

Here is a nest of large magnetite crystals.

Nest of magnetite
Nest of large magnetite crystals from Bolivia.

Here's a large single crystal of magnetite.

Large single crystal of magnetite from western Australia.

Individual large crystals like this are uncommon enough to be valued by collectors. The crystal is octahedral, opaque, and strongly magnetic, easily picked up with a kitchen magnet in spite of its fairly high density.

The magnetite in the Marquenas Quartzite is probably a placer deposit, formed in a stream bed or along a beach, where heavy, inert minerals, such as magnetite, tend to be concentrated. Here's a photograph of a modern magnetite placer in the Jemez. This is on a much smaller scale than the placers preserved in the Marquenas Quartzite, but it gives the idea.

Modern magnetite placer. Near 35.734N 106.618W

Placer deposits are of considerable economic importance, because precious metals and valuable metal ores tend to be concentrated in them. Many of the most important gold deposits are placer deposits, such as this one near Fairplay, Colorado.

Gold placer
          mine near Fairplay, Colorado
Gold placer mine Near 39 13.614N 106 0.414W

The piles of rubble visible here are tailings from placer mining. The river gravels here are rich in gold from veins in the mountains further up the valley and have been exploited for over a century.

The amateur prospector panning for gold is exploiting a placer deposit. Other precious and rare earth metals, tin, and diamonds are also extracted from placer deposits in various parts of the world.


Metarhyolite of the Vadito Group has a radiometric age of 1.70 billion years, significantly younger than the Moppin Complex. There are particularly fine outcroppings in the vicinity of Ojo Caliente.

Panorama from
        northwest of Ojo Caliente
Panorama from ridge northwest of Ojo Caliente. 36 18.416N 106 3.294W

The ridge from which this panorama was photographed is northwest of the resort, and the panorama begins looking to the northeast. The ridge itself, which continues to the south, is nearly the southernmost Precambrian exposure of the Tusas Mountains. The peak to the right in the panorama is Cerro Colorado, the southernmost peak of the Tusas, underlain also by Precambrian metarhyolite. Here's a closer look at this rock.

Xenolith. From a point just southwest of the hill top of the previous panorama.

This boulder contains a patch of darker material, which is likely a xenolith. A xenolith is a bit of country rock picked up by a body of liquid magma, which does not quite melt and remains distinct from the magma. In this case, the darker color and coarser grains of the xenolith suggest it is a mafic rock, possibly from the lower crust or upper mantle.

Exposures of Vadito metarhyolite are extensive in the southern Tusas Mountains, and all are thought to have formed from metamorphosis of high-silica volcanic ash erupted around 1.70 billion years ago. The ash came from nearby continental arc volcanism.

The Tres Piedras Granitic Orthogneiss

The village of Tres Piedras ("Three Rocks") is the eastern gateway to the Tusas Mountains. To the east is the sagebrush plain of the Taos Plateau, while the village itself is forested with ponderosa pine. The town derives its name from three large outcrops of "granite", which record the end of the Vadito back-arc basin.

Shortly after the deposition of the Vadito Group, the back-arc basin began to close up again. This was accompanied by intrusions of felsic magma, the largest of which formed the Tres Piedras Granitic Orthogneiss, known informally as simply the Tres Piedras Granite. These intrusions cut across both the Moppin Complex and the Vadito Group. Such intrusions along suture zones, which help join continental plates together, are sometimes described as stitching plutons. (Pluton is a general term for a large body of intrusive rock, derived from Pluto, the Greek god of the underworld.)

Part of the contact between the Tres Piedras Granite and the Moppin Complex is exposed in a road cut just west of the village.

Contact between Moppin Series and Tres Piedras Granite
Contact between Moppin Complex (left) and Tres Piedras Granite (right). 36 39.155N 10558.697W

Here are some close ups of rocks from either side of the contact.

Moppin schist
Moppin Complex schist
Tres Piedras granite
Tres Piedras granite

Under the loupe, the schist appears to be mostly black amphibole with a scattering of white feldspar and an occasional dark garnet. The granite is a mixture of quartz and feldspar with the occasional flake of mica.

Another, smaller, intrusion is exposed at Tusas Mountain. The age of this rock is controversial, but the best recent measurement gives the age as about 1.693 billion years. 

The Hondo Group

The other major group of formations south of the Spring Creek Shear Zone are the Hondo Group. These are all less than 1.7 billion years old and are interpreted as supracrustal rocks of the Yavapai-Mazatzal transition zone. Supracrustal rocks are rocks deposited on an existing basement.  Most of the exposures consist of a very clean quartzite called the Ortega Quartzite, which was laid down towards the end of the Yavapai Orogeny on a gently sloping surface eroded out of the older Moppin Complex and Vadito Group beds.

The Ortega Quartzite

The Ortega Quartzite originally consisted of cemented grains of almost pure quartz. It has since undergone metamorphosis to a very tough rock called quartzite. Quartzite is composed of almost pure quartz crystals packed densely together, which differs from sandstone, which consists of individual rounded grains of quartz with considerable pore space (which is sometimes filled with other minerals). Quartzite forms because quartz under stress is slightly more soluble than quartz that is not under stress. When sandstone comes under great pressure, the points where the individual grains touch take up the stress, which causes the quartz to dissolve away at the contacts. It is redeposited where there is less stress -- in the pore spaces between the original grains.

Sandstone to quartzite

Extensive outcroppings of Ortega Quartzite are found in the Tusas and Sangre de Cristo Mountains.

Kiowa Mountain in the Tusas Range is underlain by Ortega Quartzite.

Kiowa Mountain
Kiowa Mountain, underlain by Precambrian Ortega Quartzite. Viewed from 36 38.483N 106 3.481W

By a stunning coincidence, Ortega Quartzite also crops out throughout the Ortega Mountains of the southwestern Tusas.  Here is one outcropping of particularly clean muscovitic quartzite.


Muscovite quartzite. 36 25.682N 106 0.735W

Under the loupe, this stuff does indeed look like almost pure quartz, with just a few flakes of light mica. This is what is meant by "clean" quartzite.

Ortega Quartzite is also found in the Picuris Mountains to the east, where it is somewhat different in color but equally clean.

Ortega Quartzite
          in the PIcuris Mountains

          Quartzite sample from the Dixon area
Ortega Quartzite. 36 12.210N 105 54.545W

The contact between the Vadito Group and the Hondo Group is spectacularly exposed in the Pilar Cliffs in the gorge of the Rio Grande.

Pilar Cliffs
PIlar Cliffs. 36 15.738N 105 48.262W

The near-vertical cliffs are exposures of quartzose schist of the Glenwoody Formation,Vadito Group. The cliffs are capped with Ortega Quartzite. The contact between the two, which is difficult to make out in this photograph but is very distinct at close range, shows features consistent with a ductile shear zone. The beds of the Glenwoody Formation just below the contact have a pink coloration and are anomalously rich in manganese.

The very clean character of the Ortega Quartzite has been a puzzle for geologists. The formation of sandstone of nearly pure quartz in the Phanerozoic Eon almost always required repeated cycles of erosion and deposition, but the Ortega Quartzite seems to have been a first-cycle quartz unit, formed from sediments eroded directly from igneous rock. Some geologists have suggested that the unique chemistry of the ocean and atmosphere 1.6 billion years ago, when oxygen was just beginning to replace carbon dioxide and hydrogen sulfide, may account for the rapid chemical weathering of the sediments. Others point to the lack of any vegetation other than microbial mats or to extreme environmental conditions of wind and tides.

Although the Ortega Quartzite is almost pure quartz, it does contain some magnetite, and there are also occasional thin beds of what must once have been clay. These were so severely metamorphosed that they recrystallized as the high temperature polymorph of aluminum silicate, sillimanite.

Sillimate in
          Ortega Quartzite
Ortega Quartzite with sillimanite vein. 36 15.776N 105 47.610W

Sillimate in Ortega Quartzite
Ortega Quartzite with sillimanite vein. 36 15.776N 105 47.610W

Compared with the kyanite shown earlier, the sillimanite is coarser and more translucent. It also lacks the bluish coloration.

The Rinconada Formation

Although the Ortega Quartzite dominates the Hondo Group in the Tusas Mountains, other Hondo Group formations are present in the Picuris Mountains. The Rinconada Formation is probably slightly younger than the Ortega Quartzite. It includes some impressive muscovite schists.

Rinconada schist

          schist sample
Rinconada Formation. 36 19.646N 106 03.324W

The dark grains are probably crystals of staurolite or garnet, aluminum-rich minerals that are often found with muscovite in aluminum-rich metamorphic rock that recrystallized at a temperature around 600C (1100 F) at a depth of around 20 km (12 miles). The Rinconada Formation was probably originally mudstone, rich in clay, which was metamorphosed into quartz-muscovite schist. This is a rock rich in muscovite mica, which forms parallel layers along which the rock is fairly easily split.

The transition between the Ortega Quartzite and the Rinconada Formation is another ductile zone.

Rinconada - Ortega
Contact of Ortega Quartzite and Rinconada Formation. 36.201702N 105.907231W

The Rinconada Formation a hundred feet to the east (left) is a pale gray color,  but fairly abruptly changes to red without any other obvious lithological change. The lowermost beds show evidence of ductile flow, and just left of center is the contact with the Ortega Quartzite. This becomes its normal pale color to the right, but is stained bright red immediately right of the contact. The staining here looks like hematite rather than manganese minerals.

The Rinconada Formation is famed for its staurolite beds.

Rinconada Formation. 36 15.738N 105 48.262W

Rinconada Formation
Rinconada Formation. 36 15.738N 105 48.262W

The large crystals visible on the sample are staurolite crystals, along with smaller grains of garnet. Staurolite has the composition Fe2Al9O6(SiO4)4(O,OH)2. That is, it is a very aluminum-rich mineral which also contains some silica and reduced iron. Highly weathered clay is rich in alumina (aluminum oxide), and the high aluminum content of staurolite reflects the high clay content of the sediments from which this rock formed. There must also have been a modest amount of iron in the clay.

Staurolite is an index mineral, characteristic of a particular metamorphic environment. The geologist George Barrow first identified distinct zones of increasingly highly metamorphosed mudstone in the Scottish Highlands, now called the Barrovian zones. Each of these zones is marked by the appearance of a new index mineral. The earliest stage of metamorphosis produces chlorite, a mineral similar to mica. Further metamorphosis at increasing pressure and temperature produces biotite mica, then garnet, then staurolite. Later comes kyanite and sillimanite, the last occurring as the rock approaches the melting point at great depth.

Rinconada Formation
Rinconada Formation. 36 15.738N 105 48.262W

Individual staurolite crystals easily weather out of the less resistant matrix, and these are often "twinned" crystals like the ones shown above. Most twins are pairs of rodlike crystals that interpenetrate at an angle of about 60 degrees. Rare, and more valuable, are twins at right angles, which form so-called "fairy crosses." There is a single imperfect fairy cross in these samples, third from the left in the top row.

Staurolite is a fairly common mineral, but crystals this well formed are unusual. The staurolite zone represents a narrow range of temperature and pressure, near 580 Centigrade at a depth of about 28 km (17 miles). Outside this zone, staurolite is unstable.  A difference of 15 degrees Centigrade in either direction prevents staurolite from crystallizing. In addition, only very aluminum-rich sediments produce staurolite, even when they are in the right temperature and pressure zone. Anything less aluminum-rich produces only garnet. So the Pilar staurolite beds are pretty unusual.

You may be wondering why, if staurolite is unstable outside a narrow temperature range, we see any in surface rocks. The answer is that both the formation and the destruction of mineral crystals in metamorphic rock is slow even at high temperature, and very slow indeed at low temperature. If the rock is brought rapidly to the surface by tectonic forces, so that the rock is rapidly cooled, unstable minerals can survive for a very long time.

The Pilar Phyllite

A particularly interesting formation is the Pilar Phyllite, which was probably laid down just after the Rinconada Formation and is sometimes assigned to that formation.

Pilar Phyllite in
          the PIcuris Mountains

Pilar Phyllite. 36 12.661N 105 49.761W

This rock is particularly interesting because of its high content of graphite, which is thoroughly disseminated through the quartz and muscovite making up the bulk of the rock. It is unlikely we are looking at abiogenic carbon in this rock, which therefore shows the earliest signature of life in northern New Mexico. This rock was probably laid down in a shallow sea, rich with nutrients from nearby volcanic activity, in which cyanobacteria thrived and extracted carbon from the atmosphere.

Having ventured north for clues to wider events during the Mesoproterozoic, we now examine the oldest rocks of the Jemez region itself.

The northern Sierra Nacimiento Mountains

Map of the Jemez highlighting Yavapai Province
Relief map of the Jemez with Yavapai outcroppings highlighted in red.

At the westernmost edge of the Jemez is a range of very old mountains that runs almost directly north and south. This is the Sierra Nacimiento. To its west is the San Juan Basin of the Colorado Plateau, and moisture picked up by the prevailing winds as they cross this long stretch of flat ground is wrung out over the Sierra Nacimiento. The rainfall supports lush forests of ponderosa pine and other conifers that mantle peaks and valleys softened by erosion. Between the Sierra Nacimiento and the Jemez Plateau to its east is the valley of the Rio Guadelupe and Rio de las Vacas. This is a favorite camping area for weekend visitors, but also supports small herds of cattle and  some logging. One may even encounter Hispanic cowboys driving their cattle along the few highways in the area, as they have done for over three centuries.

Senorita Canyon cattle drive

Cow drive. 35 59.505N 106 51.981W

Señorita Canyon cuts across the Sierra Nacimiento east of Cuba, separating the northern third of the range from the remainder of the mountains. The northern Sierra Nacimiento, also known as the San Pedro Mountains, is an area of parks (mountain valleys) and gentle hills forming a high plateau that reaches to a maximum altitude of 3232m (10,605').

Though part of the Jemez Mountains, the Sierra Nacimiento has a very different origin and history from the rest of the region. The bulk of the Sierra Nacimiento Mountains is composed of Precambrian rocks that have been repeatedly thrown up by tectonic forces during the Phanerozoic Eon. In the northern Sierra Nacimiento, these include the oldest rocks in the Jemez region.


Map of the Jemez highlighting metamorphic outcrops
Relief map of the Jemez with San Pedro metamorphic outcroppings highlighted in red.

The San Pedro Mountains are a patchwork of plutons, most of which have never been radiometrically dated. There are only isolated remnants of the older metamorphosed sedimentary and volcanic beds that the plutons intruded, and the relative ages of the plutons must be inferred from the contact relationships between them. Such radiometric ages as we have suggest that the older metamorphic rocks are correlated with the youngest Moppin Complex or oldest Vadito Group beds.

These rocks are exposed in Rito de las Perchas, a pleasant mountain park (grassy valley) in the San Pedro Parks Wilderness accessible by trail. The bulk of the metamorphic beds here appear to be a metarhyolite not unlike the Vadito metarhyolite. The area is gentle rolling hills and exposures tend to be poor.



San Pedro metavolcanics in Rito de las Perchas. 36.065763N 106.796216W

This was likely an ash flow from a volcanic arc associated with the Yavapai Orogeny, since recrystallized under heat and pressure. Samples of this formation were radiometrically dated in 1974 as 1.8 billion years old, but that date was measured long enough ago that it is not terribly reliable.


Map of the Jemez highlighting tonalite outcrops
Relief map of the Jemez with San Pedro tonalite outcroppings highlighted in red.

One of the older plutons is composed of a rock variously described as tonalite, graniodiorite, or quartz diorite. All imply a rock with moderate to high quartz content and much more plagioclase feldspar than alkali feldspar. According to the geologic map, this is exposed along the Rito de las Perchas and in the road cut on the forest road to the south.


San Pedro tonalite along Rito de las Perchas. 36.0554323N 106.808372W



San Pedro tonalite in road cut. 36.026875N 106.796381W

The sample is unusually rich in dark iron minerals, mostly biotite. These have produced some staining, but the feldspar is white plagioclase rather than pink potassium feldspar. There is also abundant muscovite. The rock is not actually deficient in potassium, but is rich in aluminum, which has tied up the potassium as micas rather than potassium feldspar.This hints at an origin from melted clay-rich sedimentary rock. Such granites are classified as S-type granites by geologists.

The San Pedro Quartz Monzonite

Map of the Jemez highlighting San Pedro Quartz
Relief map of the Jemez with San Pedro Quartz Monzonite outcroppings highlighted in red.

Rapakivi quartz monzonite

Rapakivi quartz monzonite. 36 01.394N 106 51.005W

Quartz monzonite is an intrusive rock containing roughly equal quantities of alkali and plagioclase feldspar and between five and twenty percent quartz. Like the Maquinita Granodiorite, the San Pedro Quartz Monzonite is considered a calk-alkali rock and, with an age of 1.73 billion years, it is similar in age to the Maquinita Granodiorite.

This outcrop has a small shear zone crossing it.

Shear zone

This is a zone in which the rock has been deformed, forcing the crystals to align into bands. The rock to the south (right) shows no signs of deformation. That to the north (left) gradually transitions from the sheared appearance in the center of the photograph to an undisturbed texture. This is a much smaller, more local equivalent of the Spring Creek Shear Zone further north. Such shear zones are associated with deep crustal movement, with the rock on opposite sides moving past each other. This resembles a fault, but occurs at such great depth that the rock flows rather than fractures. In other words, this occurs at depths below the brittle-ductile transition. Similar shear trending east to east-northeast is apparently common, judging from the geologic paper that first described this formation.

Nearby is a big patch of more mafic rock.


This is probably not a dike; it’s too localized. It looks like a big blob of country rock that broke off and sank into the body of magma.

An outcrop to the south has clear rapakivi texture.

          quartz monzonite

Rapakivi texture. 36 01.354N 106 50.945W

The weathering of this surface helped bring out the texture Here’s a sample.

Rapakivi quartz monzonite texture

The rapakivi texture refers to the big pink rounded crystals. These are orthoclase with a rim of oligoclase, a sodium-rich plagioclase feldspar. You can see the rim particularly well on the big crystal towards the right side of the sample. The presence of big, fat, and happy orthoclase crystals with reaction rims makes this a rapakivi quartz monzonite; the oligoclase in the rims makes it a particular kind of rapakivi called vyborgite. The bluish grains of quartz are quite striking. The rock also contains dark grains of biotite.

Though the presence of large orthoclase crystals is quite common in silica-rich intrusive rock, true rapakivi texture is fairly rare. It is usually associated with so-called A-type granite, of which we'll learn more shortly.

Further north, the monzonite shows a system of joints trending northwest to southeast, cutting across the shear zones noted earlier.

          quartz monzonite

The joints are oriented nearly vertically.

          quartz monzonite

Jointing in the San Pedro Quartz Monzonite. 36.0270413N 106.849181W

The joints likely formed when this rock was at a shallower depth than when the shear zones were formed. They may be related to the most recent upthrust of this area, less than 25 million years ago.


Biotite is a mica with the formula KFe3AlSi3O10(OH)2. It is common for magnesium to substitute for some of the iron and fluorine to substitute for some of the hydroxyl. (Fluoride and hydroxyl ions have the same charge and very nearly the same diameter.) Biotite is very common in igneous rocks, being present in everything from extrusive, low-silica basalt to intrusive, high-silica granite.

Biotite differs from muscovite in having iron hydroxide (or fluoride) rather than aluminum hydroxide act to bond pairs of phyllosilicate sheets together. This modifies the structure slightly. The ferrous hydroxide sheet does not consist of open rings of cations bonded by hydroxyl, as with the aluminum hydroxide sheet in muscovite, but of a dense sheet of packed iron cations bonded by hydroxyl. Instead of two aluminum ions joining each pair of rings of the two phyllosilicate sheets, three iron or magnesium ions join each pair of rings. A mica having two metal ions joining each pair of rings is described as dioctahedral, while one having three metal ions join each pair of rings is described as trioctahedral. The octahedral refers to a site in the middle sheet surrounded by six oxygen or hydroxyl ions where a metal ion of the right size could potentially be located.

Here's a sample of particularly coarsely crystallized biotite.


Biotite crystals in a sample of granite.

Like muscovite, biotite has a single perfect cleavage plane that allows the mineral to be split into very thin elastic sheets. However, biotite can usually be distinguished from muscovite by its very dark color and higher density.

Muscovite-biotite granite

Map of the Jemez highlighting muscovite biotite
Relief map of the Jemez with muscovite-biotite outcroppings highlighted in red.

The San Pedro Peaks area is underlain by a pluton that appears to be younger than the San Pedro Quartz Monzonite. This also crops out in the upper canyon of the Rio Puerco and Rito Resumidero.

          biotite granite
Muscovite-biotite granite at Rito Resumidero falls. 35 35.946N 106 53.533W

Muscovite biotite granite
Muscovitebiotite granite at Rito Resumidero falls. 35 35.946N 106 53.533W

This is described a muscovite biotite granite, containing microcline and quartz and smaller quantities of sodium-rich plagioclase and muscovite and biotite. Such two-mica granites are thought to form primarily from aluminum-rich sedimentary beds heated by basalt underplating.


The youngest sizable pluton in the San Pedro Mountains is thought to be a mass of leucogranite, which is a light-colored granite containing very little dark mafic minerals.

Leucogranite near Deer Mountain. 36.0162173N 106.7976472W

Leucogranite near Deer Mountain. 36.0162173N 106.7976472W

Leucogranites are usually interpreted as a product of the melting of thickened crust composed almost entirely of clay minerals.The leucogranite here is relatively fine grained as well as almost devoid of any iron minerals except traces of biotite. In some locations, it has bluish quartz and plagioclase grains thought to have melted out of the surrounding quartz monzonite.

With no radiometric date, we have almost no age constraints on this pluton. We know only that it is younger than 1.73 billion years and older than 25 million years, the age of a patch of sedimentary rock overlying it to the north. But it is though to be Precambrian, and I would hazard a wild guess it is around 1.4 billion years old, the age of similar granite intruded further south that I'll describe shortly.

The southern Sierra Nacimiento Mountains

South of Señorita Canyon, the Sierra Nacimiento rises to maximum elevations of 2424m (9264') at Big Mountain, 2887m (9471') at San Miguel Mountain and 2750m (9022') at Pajarito Peak.

Guadalupe Box
Pajarito Peak. Looking north from 35 35.946N 106 53.533W

The mountains here are less lushly forested than further north, with pinon scrub dominating the lower slopes. Much of the southernmost Sierra Nacimiento belongs to the Zia and Jemez Pueblos, while most of the rest is National Forest.

Big Mountain is underlain by Phanerozoic rocks, but San Miguel Mountain and Pajarito Peak are underlain by Precambrian rocks.The Precambrian rocks of the southern Sierra Nacimiento Mountains include some that are the right age and character to be assigned to the Yavapai-Mazazatl transitional zone basement.

Map of the Jemez highlighting Mazazatl Province
Relief map of the Jemez with Mazatzal province outcroppings highlighted in red.

The San Miguel Gneiss

The central Sierra Nacimiento is dominated by the San Miguel Gneiss, which has a radiometric age of 1.695 billion years. This is consistent with this rock being emplaced during the Yavapai Orogeny.

San Miguel
          Gneiss outcropping

          Miguel Gneiss sample
San Miguel Gneiss. 35 50.677N 106 51.340W

The minerals here are granitic: biotite, quartz, and alkali feldspar. The biotite forms distinct bands in the rock, leaving no doubt it is a gneiss. Its composition is that of a monzogranite, with more plagioclase than alkali feldspar and around 25 percent quartz. Its age and character somewhat resemble the Tres Piedras Granitic Orthogneiss, and it may have formed by the same process along the suture zone.

Intrusive formations that have become exposed at the surface sometimes retain beds of the overlying country rock that have not quite eroded away. These are known as roof pendants and take the form of localized outcroppings of rock that are different in character from the rock forming the intrusion. They are also rootless; that is, they do not connect with any deeper structure or nearby outcroppings. There is a roof pendant in the San Miguel Gneiss that caught the attention of geologists mapping the area. It is just a couple of hundred feet across, is located on a heavily forested hillside, and is not that well exposed. I have to admire the geologists who were thorough enough in their mapping to discover it.


Hornblendite outrcop in the San Miguel Gneiss.. 35 50.530N 106 51.110W

This rock is mostly hornblende, with just a scattering of feldspar. This means a low silica content, perhaps as low as 45%. Such rocks are described as ultramafic and they are fairly uncommon. Its location on the boundary of two ancient crust provinces suggests this may be a remnant of oceanic crust trapped between the Yavapai and Mazatzal Provinces when they merged 1.6 billion years ago.

Guadelupe Box

There is a beautiful and easily accessible outcropping of Precambrian quartz monzonite at the Guadalupe Box.

Guadalupe Box
Southern entrance to Guadalupe Box. 35 43.876N 106 45.687W

The road here passes through a pair of tunnels, the Gilman Tunnels, which were blasted out the rock in the 1920s to make way for the Santa Fe North Western Railroad. By the time the tunnels were completed, in August 1924, they had accounted for half the cost of construction of the railroad. Timber was loaded onto the rail cars at Deer Creek Landing and ruins of logging camps can still be found on Peggy Mesa, west of the tunnels. The tunnels and settlement were named for William H. Gilman, the vice president of operations of the railroad.

The Great Depression and a series of railroad accidents brought logging to a near halt by 1937, but logging resumed under the New Mexico Timber Company in the 1940s, which constructed the town of Gilman, Today this is a small cluster of homes with a craft shop. The company also removed the rails to make way for trucks, which were both cheaper and safer to operate. Logging again came to a near halt in the 1960s and the tunnels were deeded to the Forest Service, which renovated the bridges approaching the tunnels and paved the road for automobiles.

The Guadalupe Box is a beautiful area; I'll let these next pictures speak for themselves.

Guadalupe Box
35 44.045N 106 45.895W

Guadalupe Box

Guadalupe Box
Guadalupe Box. 35 44.106N 106 45.898W

The quartz monzonite itself is not much to look at when weathered.

Weathered Guadalupe Box granite

But fresh surfaces reveal just how gorgeous a stone this is.

        Box granite

Guadalupe Box granite
Fresh surface of Guadalupe Box granite

Strictly speaking, this beautiful rock is a hornblende-biotite quartz monzonitic gneiss. The large feldspar crystals are more typical of high-silica intrusive rocks than the rapakivi texture of the San Pedro Quartz Monzonite: They are not as rounded and do not show an oligoclase rim. The quartz is less striking and this rock is more abundant in iron-rich minerals. It has also been metamorphosed, though the foliation from metamorphism is barely evident here.

The hornblende in this rock is a very common example of an amphibole mineral.


Amphiboles are members of the family of silicate minerals called inosilicates, which are characterized by long chains of silica tetahedra. Amphiboles are those inosilicates in which the each silica tetrahedron is joined to two or three neighbors so that the silica backbone consists of two parallel chains of tetrahedra joined together:

Sorosilicate double chain

As with other silicate minerals, it is possible for aluminum to substitute for some of the silicon. Even with no aluminum substituted in the chain, additional metal ions are required to balance the negative charge of the backbone. As with mica, these are accompanied by hydroxyl groups. Pairs of double chains face each other, with the apical oxygens on the inside bonded to a strip of metal ions. Each such combination of two double chains bonded by metal ions looks a little like an I-beam in cross-section. The "I-beams" then interlock, with additional metal ions holding the "I-beams" in place.

End view of amphibole
Amphibole structure viewed end-on to the double silica chains.

The metal ions holding the pairs of double chains together are shown in gray, with the associated hydroxyls in green. The metal ions locking the resulting "I-beams" together are shown in yellow. The amphiboles show two cleavage planes corresponding to lines drawn through the empty spaces above and below each "I-beam".

Hornblende is a rather general term for amphiboles rich in iron. A typical formula would be Ca2Fe5Si8O22(OH)2. The grey ions in the previous diagram would then be iron and the yellow ions would be calcium. However, magnesium substitutes freely for iron, sodium substitutes for calcium, and aluminum can substitute for both iron and silicon, producing a wide range in compositions.

The monzonitic gneiss of Guadalupe Box is estimated to be about 1.6 billion years in age. Its age and location indicate that it belongs to the Mazatzal Province. This province contains island arc and microcontinent crust that formed around 1.7 to 1.65 billion years ago, based on isotope model ages, and was accreted to Laurentia during the Mazatzal Orogeny 1.65 to 1.60 billion years ago. This orogeny jammed the Jemez subduction zone to create the Jemez Lineament. It also deformed the crust northward into the Yavapai Province, resetting some radiometric clocks and partially obscuring the boundary between the provinces to help form the Yavapai-Mazazatl Transition Zone.

Cañon de San Diego

The only Precambrian rocks exposed within the Jemez proper, in Cañon de San Diego at Soda Dam, likely belong to the Mazatzal Province.

Cañon de San Diego is one of two natural highways into the Valles caldera and the only one that is open to the public today. (The highway routes from Cuba and Los Alamos are highly engineered.) Cañon de San Diego can be visited via State Road Four, which branches off U.S. 550, the Bernalillo – Farmington highway, at San Ysidro. The canyon is a geologist's playground, cutting through young volcanic beds of the Jemez to expose colorful Mesozoic and Paleozoic sedimentary beds of the eastern edge of the Colorado Plateau. The upper canyon is heavily forested with ponderosa pine, which gives way to pinon scrub forest in the lower canyon. Much of the lower canyon belongs to the Jemez Pueblo, while the upper canyon includes the village of Jemez Springs, several private developments, Hummingbird Music Camp, and numerous National Forest campgrounds. These make the canyon one of the more densely populated portions of the Jemez.

The ancestors of the Jemez Pueblo settled the area beginning around the 13th century, and some of their ruined pueblos can still be found in the canyon or atop the high mesas on either side. The pueblo of Giusewa was built at this time on the current site of Jemez Springs. The Spanish established a church beginning in 1621, San José de los Jémez, which can be visited today as part of the Jemez Historic Site.

San José de los
San José de los Jémez. 35 46.700N 106 41.227W

However, the church was abandoned by 1640, and the pueblo was abandoned during the Pueblo Revolt of 1680 in favor of pueblos in more defensible locations. The inhabitants never returned, ultimately settling further down canyon where their descendants live today.

San Diego Canyon roughly coincides with a major fault zone, the Jemez Fault Zone, which is part of the western margin of the Rio Grande Rift. At Soda Dam, the fault has brought up some of the Precambrian basement rock. This is a granite gneiss with a radiometric age of about 1.6 billion years.

Precambrian outcropping at Soda Dam

View to the north from south of Soda Dam, which is partially visible as the low ridge just across the road. The Precambrian

gneiss towers over the road on the left. 35 47.481N 106 41.208W

Precambrian granite gneiss southwest of Soda Dam
Granite gneiss southwest of Soda Dam.

Granite gneiss
          hand sample
A sample of granite gneiss from Soda Dam.

As with the feldspathic schist from the Tusas, this rock is mostly quartz and feldspar with some mafic minerals. The grains in the rock are oriented in a common direction, which is not true of an ordinary granite, whose crystals are oriented at random. This indicates that the rock was once under great heat and pressure, which caused it to recrystallize parallel to the shear forces it was experiencing. In the case of the granite gneiss of the Soda Dam area, the recrystallization is fairly subtle and the rock only mildly metamorphosed, so that it superficially still resembles fine-grained ordinary granite. Strongly metamorphosed granite gneiss shows a characteristic banding that is not evident at Soda Dam, but which we saw earlier with the San Miguel Gneiss. The lack of schistose layering indicates this rock lacks mica and suggests that it is not highly enriched in aluminum.

This relatively small outcropping of Precambrian gneiss has not been correlated with nearby Precambrian formations. The nearest Precambrian rocks that have been assigned formation names are in the Sierra Nacimiento Mountains just to the west. The rock here doesn't fully match the description of any of the granites further west, though its age is similar to the older of these granites.

The anorogenic granites of 1.4 billion years ago

The period between 1.8 and 0.8 billion years ago, from the late Paleoproterozoic to the mid-Neoproterozoic, has been dubbed the "Boring Billion" by some geologists. This was a long period of relative tectonic stability, in which the atmospheric oxygen level held fairly steady at less than 10% of its current value, and in which life evolved only slowly. However, this time period was not completely uneventful. For one thing, sexual reproduction first appeared around 1.2 billion years ago. By 1.4 billion years ago, the supercontinent of Columbia was showing signs of breaking up, only to reassemble almost at once into Pannotia. And something was cooking under what is now western North America.

Map of the Jemez highlighting Joaquin Granite 
Relief map of the Jemez with Joaquin Granite outcroppings highlighted in red.

Following the end of the Mazatzal Orogeny 1.6 billion years ago, the Jemez region experienced a period of tectonic quiescence. This ended around 1.4 billion years ago, when the Precambrian rocks of the Mazatzal and Yavapai Provinces were intruded in many locations by huge granite or granite-like bodies, called batholiths. The 1.4 billion year batholiths are found across the western United States, constituting fully 15% to 40% of the Precambrian surface. These batholiths point to a major episode of widespread crustal heating whose cause is still hotly debated by geologists.

Granite from one such batholith can be found in the Rio Guadalupe Canyon, at the mouth of the tributary Joaquin Canyon.

Joaquin Granite

Joaquin Granite

Joaquin Granite
Joaquin Granite. 35 46.421N 106 47.519W

This is the Joaquin Granite, which is the most common Precambrian rock in the southern Sierra Nacimientos. It's a true granite with a radiometric age of 1.424 billion years.

Another outcropping of the Joaquin Granite is found further north in a road cut.

Joaquin Granite. 35 49.459N 106 50.160W

The Joaquin Granite obscures the precise boundary between the Yavapai and Mazazatl provinces, since it was intruded close to that boundary long after both provinces had accreted to Laurentia.

Across the Rio Grande Rift from the Sierra Nacimientos and the Jemez Mountains is the southern Sangre de Cristo Mountains. This area is part of the Mazatzal Province, and the rocks are a confusion of 1.6 to 1.7 billion year old biotite schist and granite gneiss intruded by 1.4 billion year old granite of the anorogenic event.

        intruding greenstone along the Hyde Park road
Granite intruding greenstone along the Hyde Park Road northeast of Santa Fe. Near 35 42.774N 105 53.670W

Further up the road, the entire assemblage is distorted from intense metamorphism.

Hyde Park Road heavily
        metamorphosed outcropping
Heavily deformed rock further east along Hyde Park Road

The biotite schist is typical of metamorphosed oceanic crust, and is evidence that much of the Mazatzal Province to which it belongs came from ancient island arcs that accreted onto the southern coast of Laurentia.  

The 1.4-billion-year-old batholiths found throughout the Yavapai and Mazatzal Provinces have been described as anorogenic, meaning that they do not appear to have been associated with a major episode of mountain building, such as from a collision of two continental plates. However, there is some debate about this, with a few geologists claiming that there is evidence for mountain building at this time in the Picuris Mountains. Another episode of accretion occurred along southern Laurentia from 1.45 to 1.3 billion years ago to form the Granite-Rhyolite Province, and while the Jemez region was far inland from the continental margin by this time, the coincidence in timing suggests a connection. Possibly the continuing subduction produced heating far inland. So geologists have hedged their bets: These granites are described as A-type granites, with the A standing for anorogenic; but it can also stand for alkaline, without any judgment on its mode of origin. The A-type granites have a distinctive chemical composition, being rich in silica and alkaline metals (sodium and potassium) and having a high ratio of iron to magnesium and a low calcium content.

Whatever the cause of the crustal heating that produced these batholiths, it had the effect of converting the mafic crust that originally assembled into the Yavapai and Mazatzal Provinces to mature crust. The average composition of the crust changed from that of basalt to that of andesite, as the more mafic material settled to the base of the crust and delaminated into the mantle.

Dikes and pegmatites

The heating event 1.4 billion years ago left its mark on the Tusas Mountains as well. Here the younger rock mostly takes the form of dikes. For example, dikes of almost pure quartz cut across the Moppin Complex on Hopewell Ridge.

Quartz vein
Large quartz vein. Near 36 38.339N 106 07.740W

Dikes form when magma forces its way through a fissure in the country rock and then cools in place. They are perhaps the most common form of intrusive body or pluton.

An impressive pair of pegmatite dikes are found at a road cut in the southernmost Tusas Mountains.


Pegmatite dike

Pegmatite dike
Pegmatite dikes. 36 24.192N 106 01.266W

These dikes are located just a few yards from each other along the same road cut. It is quite common to see multiple dikes running parallel with each other. Where a large number of parallel dikes are found in a region, geologists often refer to them as a dike swarm. There are numerous pegmatite dikes in the Tusas Mountains, though probably not so many that they constitute a dike swarm.

Notice that the foliation in the country rock appears to be present in the pegmatite as well. This suggests that there was a significant episode of metamorphic deformation following the emplacement of the pegmatites.

Pegmatites are notable for the presence of very coarse crystals, sometimes of quite unusual minerals.

Close up of

Closer up of
Pegmatite dikes

This pegmatite is full of large crystals of quartz, feldspar, muscovite, and accessory minerals. Such large crystals do not form from slow cooling alone. The magma from which they form must also be rich in  water vapor, which greatly lowers its viscosity. This water vapor also accounts for the presence in pegmatites of minerals such as mica that include water in their crystal structure.

Pegmatites are thought to form from the very last part of a granitic magma chamber to crystallize, and so they tend to contain unusual minerals containing incompatible elements. Incompatible elements are elements having a combination of ionic radius and electrical charge that is significantly different from those of the more common rock-forming elements. As a result, these elements are reluctant to enter ordinary rock-forming minerals as a trace constituent. For example, manganese is nearly identical to iron in its charge and ionic radius, making it a compatible element, and it commonly substitutes for iron in iron-bearing minerals. This is why distinctive manganese minerals are not terribly common, even though manganese is a fairly abundant element. Boron, on the other hand, has a charge and radius unlike the more common elements, making it an incompatible element. Though a very rare element, it tends to concentrate in the residual magma fluids that form pegmatites, which therefore often contain tourmaline or other distinctive boron minerals. Other incompatible elements found in pegmatites include lithium, beryllium, fluorine, tin, niobium, tantalum, and certain lanthanide metals. The unusual composition makes pegmatites attractive to prospectors, and there are many old mines in the Tusas Mountains.

One such mine is the Joseph Mine, located just a couple of miles north of the small resort of Ojo Caliente. We saw a panorama of the hilly terrain north and west of the resort earlier in this chapter. This terrain is underlain mostly by Vadito Group metarhyolite, amphibolite, and schist, intruded in numerous locations by pegmatite dikes. The Joseph Mine itself is located in a large pegmatite plug that has intruded the boundary between metarhyolite and amphibolite outcrops. The mine takes the form of a sizable open pit.

Joseph Mine viewed
          from the west

Panorama of Joseph Mine

Joseph Mine. 36 19.646N 106 03.324W

The A-type pegmatites of the Tusas Mountains are rich in aluminum, and the combination of high aluminum and potassium content is favorable for forming muscovite mica. Mica has been extensively mined from the Tusas Mountains and was the principal product of the Joseph Mine. Some of the mica here was truly spectacular, forming “books” (individual crystals) exceeding three feet in diameter. Even today, it is easy to find mica books six inches across.

Mica books at Joseph Mine

Mica books at Joseph Mine. Car keys at lower right for scale.

More such books are exposed in the short adits (horizontal tunnels) cut into the pegmatite nearby. These are difficult to extract intact, but I managed the following specimen.

Large muscovite book

Muscovite from Joseph Mine

Unfortunately, this fell apart before I could wrap it up. But I managed to get some other large specimens home intact.


Muscovite from Joseph Mine

It’s not obvious in this photograph, but this sample is nearly five centimeters (two inches) thick. It's a single crystal of muscovite.


Amphibolite that is intruded by pegmatite often contains almandine garnet near the contact. Individual garnet crystals can be found weathered out of the contact between the pegmatite and the adjoining amphibolite of the Joseph Mine. Few of these are gem quality, but they can still be fun to hunt down and collect. As with fossil hunting, you have to train your eye to spot garnets mingled with the pebbles along the slopes below the amphibolite.


Garnets from Joseph Mine

The crystals are imperfect, but you can see crystal faces on the better samples.

Garnet is an example of a nesosilicate, in which isolated silica tetrahedra are completely surrounded by metal ions that provide charge balance. The composition of garnet is highly variable; almandine typically has the composition Fe3Al2(SiO4)3, but almost any metal ion with a charge of +2 can substitute for the iron and either ferric iron or chromium can substitute for the aluminum.

Garnet is found almost exclusively in high grade aluminum-rich metamorphic rock. At the Joseph Mine, the metamorphism was caused by a nearby hot body of magma (the pegmatite), a process which is called contact metamorphism. This is in contrast with the regional metamorphism that produced the extensive metamorphic formations of the Tusas and Picuris Mountains, which was caused by an orogeny.

Staurolite beds almost always contain garnets as well. The reverse is not true: Most garnet-bearing rock does not contain staurolite. Garnets do not require as unusually high an aluminum content as staurolite, and they are stable over a much broader range of pressures and temperatures, so that garnet is a fairly common mineral in aluminum-rich metamorphic rock.


One of the unusual minerals found at the Joseph Mine is tourmaline. This is found mostly on the western rim of the mine, close to the contact of the pegmatite with the host amphibolite.

Tourmaline beds
          at Joseph Mine

Tourmaline samples

Tourmaline is a cyclosilicate mineral, whose basic framework is stacked rings of silica tetrahedra. These are bonded to triangular borate ions by various metal ions. The type of tourmaline found at Joseph Mine is schorl, in which the metal ions are predominantly iron and sodium, yielding a formula NaFe3Al6(BO3)3Si6O18(OH)4. You can see that the schorl takes the form of long black striated rods. A wide range of metal ions can substitute for sodium, iron, aluminum, and silicon, making tourmaline one of the minerals most variable in composition. The sodium and borate come from the pegmatite, while the iron comes from the amphibolite, and silica and aluminum comes from both. Alteration in the composition of country rock by fluids from an intrusion is called metasomatism

Along with muscovite and accessory minerals like garnet and tourmaline, the pegmatite at Joseph Mine is composed of quartz and alkali feldspar. These are visible in this outcrop.

Feldspar in

The reddish mineral is probably microcline, while the white could be either albite or quartz. There is also considerable muscovite.

I picked up a nice sample of feldspar here.


This is a cleavage fragment from a single large crystal. One can hold the sample to the light and see that the entire surface reflects the light at the same angle. The fine striations suggest that this is perthite, composed of thin alternating layers of albite and microcline, which separate from each other as the feldspar slowly cools.

Pegmatite dikes are also found in the San Miguel Gneiss This example is poorly exposed, but gives some idea.

          exposed pegmatite in San Miguel Gneiss

          Miguel pegmatite
Pegmatite in the San Miguel Gneiss.. 35 50.581N 106 51.543W

This pegmatite is rich in quartz with some feldspar, but little of any other kind of mineral. Pegmatites that contain few accessory minerals are known as simple pegmatites. It might also be classified as a leucogranite. Some quite sizable leucogranite outcrops occur in the southern Sierra Nacimientos.

Nearby is another dike of very different character.

Mafic dike

San Miguel mafic dike
Mafic outcrop in the San Miguel Gneiss.. 35 50.574N 106 51.527W

This appears to be an ultramafic intrusion of some kind. Under the loupe, it looks a little like the hornblendite from earlier in this chapter, but with only the barest scattering of feldspar and with some significant content of biotite. My geologic map for this area notes that there are localized outcrops of schistose amphibolite in the San Miguel Gneiss, though it does not specifically show this one on the map. But that's probably what this is. Was this a mafic dike, since heavily metamorphosed, or another bit of ocean crust trapped when the Mazatzal and Yavapai Provinces were sutured together? And is there any significance to its location right next to a pegmatite dike?

While it is a common view that pegmatites form from the last fraction of a magma to solidify, it is also possible that some pegmatites form from the reverse process, where regional metamorphism heats rock in the middle crust just enough for the rock to begin to melt. This melt will be rich in volatiles and incompatible elements, and if it then moves from its source region to a higher level of the crust, it could produce a pegmatite difficult to distinguish from one formed from the last liquid fraction of a large magma body. The chief difficulty with this theory is explaining how the small amount of melt separates from its source rock. Melted rock is expected to cling to the remaining solid rock like water in a sponge. None can be extracted until the source rock has melted enough for its pore spaces to be saturated with magma.

Another kind of mafic dike intrudes an outcrop of the San Pedro Quartz Monzonite along State Road 126 in the northern Sierra Nacimiento Mountains.

Lamprophyre dike

Possible lamprophyre dike in San Pedro Quartz Monzonite. 35 59.672N 106 49.325W

This dike is prominent enough to be shown on the geologic map for this area. A close examination shows a feature not often seen in mafic dikes.


There are rather large crystals of orthoclase in the otherwise fine-grained dike rock. These somewhat resemble the distinctive orthoclase crystals of the nearby rapakivi quartz monzonite. A sample:


Furthermore, while washing the sample to prepare it for its portrait, I realized that there are bluish quartz grains in the rock that are elongated in one direction. (You can see one just above and to the left of the center of the sample.) These features suggest this may be a kind of lamprophyre called vogesite. The geologic map for this area indicates that lamprophyre dikes are present in the quartz monzonite. So there it is.

Lamprophyres are very low-silica, high-potassium rocks formed in small volumes by very slight melting of the earth’s upper mantle. They are characterized by porphyritic texture, including xenocrysts of feldspar and quartz. Xenocrysts are individual mineral grains that are in some way foreign to the rock, such as grains melted out of the surrounding country rock. Lamprophyres are classified according to the dominant minerals in the ground mass, and a vogesite is a lamprophyre whose ground mass is made up mostly of amphiboles and microcline. That appears to be the case here.

The silica-rich rock making up the anorogenic batholiths of the western United States could not have formed directly from silica-poor magma produced in the mantle. The primitive magma must have risen to the base of the crust because it was less dense than the upper mantle, then spread out laterally because it was more dense than the overlying crust. Only small quantities of this magma reached the surface as mafic dikes. This magma was very hot, and it provided both a source of water vapor and heat to melt the more silica-rich rock above it. This rock was already rich in water-containing minerals from its island arc origin and so was fertile for magma production. This magma further differentiated, leaving silica-poor minerals at the base of the crust while continuing to ascend to become part of a more silica-rich upper crust. This zoning of the crust is typical of continental crust throughout the world today. It is possible that delamination subsequently removed the residual iron- and magnesium-rich rock at the base of the crust, increasing the buoyancy of the crust even further.

Almost all the Earth's continental crust is mature, yet geologists think almost all of it must have started out juvenile. One supposes that the formation of a large area of juvenile crust must somehow trigger the subsequent heating that matures the crust, perhaps by trapping heat produced by the relatively abundant radioactive minerals in the crust, but the process is still not well understood.

The following animation illustrates the formation of the crust under New Mexico. It is based on radometric dates for Precambrian rocks in the region, with a dot appearing at the location from which each sample was taken at the time corresponding to its radiometric age. This shows the emplacement of Precambrian rocks from 1800 million years ago to 1110 million years ago. Note how rocks are first formed in southern Colorado and northern New Mexico (Yavapai province) then activity abruptly spreads south (Mazatzal provice.) There is a long period of quiescence, then the anorogenic pulse of magmatism takes place at about 1.4 billion years. There is a final small burst of activity in southernmost New Mexico corresponding to the assembly of Rodinia at around 1.1 billion years.

Precambrian crust formation animation

Next chapter: When the Jemez was beachfront property.

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