The previous chapter may be found here.
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
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.
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
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.
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
|Cenozoic (66 Ma to
||The Age of Mammals
|Quaternary (2.58 Ma
||The Age of Man|
|Holocene (11.7 ka to present)
|Pleistocene (2.58 Ma to 11.7
|Neogene (23 to 2.58
|Pliocene (5.3 to 2.58 Ma)
|Miocene (23 to 5.3 Ma)
|Paleogene (66 to 23
|Oligocene (34 to 23 Ma)
|Eocene (56 to 34 Ma)
|Paleocene (66 to 56 Ma)
|Mesozoic (252 to 66
||The Age of the Dinosaurs
|Cretaceous (145 to 66
|Jurrassic (201 to 145
|Triassic (252 to 201
|Paleozoic (540 to 252
|Permian (299 to 252
|Pennsylvanian (323 to
|Mississippian (359 to
|Devonian (419 to 359
|Silurian (443 to 419
|Ordovician (485 to
|Cambrian (540 to 485
|Proterozoic (2500 to
to 560 Ma)
to 1000 Ma)
(2500 to 1600 Ma)
|Archean (4000 to 2500
|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.
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 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.
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 combined
chemically with reduced iron and sulfur dissolved in the ocean
water. 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).
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
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. The eukaryotes have cells with true nuclei and a complex internal structure. Their modern descendants include animals, plants, and fungi. 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.
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.
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. Presumably 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 at about 1.8 billion years ago. This supercontinent has been given many names: Nuna, Hudson, Protopangea, Columbia.
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 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 1.9 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
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 likely 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
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. 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 1.9 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 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).
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
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 suture where, some 1.7 billion years ago, two continental plates
collided and merged. The collision 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,
de Cristo, the Sierra
Nacimiento, and the Sandia
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. Magnetic fields are anomalously high along the Lineament. There is also 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 fossil subduction zone 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.
... these waters have been recommended by Doctor Nagle, of Santa Fe, in many chronic diseases, and always with success.
— Lieutenant Willam 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
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 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, and the rocks are classified according to their mineral content and degree of foliation. The latter is the extent to which the minerals in the metamorphic rock have segregated into distinct bands in the rock. 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.
North of Spring Creek is Hopewell Ridge, which is composed mostly of rocks assigned to the Moppin Suite. These are among the oldest rocks found in northern New Mexico, with an estimated age in excess of 1.75 billion years, and are typical of the Yavapai Province. The Moppin Suite consists of thick beds of mafic volcanic rock interbedded with occasional thinner beds of felsic volcanic rock and sediment.
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.
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 Suite rocks were erupted in a marine
environment, as part of an island arc. Another indication is the
presence of well-preserved pillow basalts, erupted under water, in
Moppin Suite outcrops in the Brazos
The Moppin Suite 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 leptite along Hopewell Ridge is an example of a felsic member of the Moppin Series.
Leptite is a metamorphic rock 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 leptite here is schistose, having a
laminated structure as shown by the thin layers of the mafic
mineral, and is sometimes described as feldspathic schist.
The minerals in leptite 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 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. A cube is an example of 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 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:
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.
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 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.
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 leptite 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 leptite 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. The thin layering is
typical of shale and may indicate that this is actually a
Another clue to the history of northern New Mexico in the
Precambrian is the presence of calc-alkaline igneous rocks
within the Moppin Suite. The most widespread in the Tusas
Mountains is the Maquinita Granodiorite, which has been dated at
1.755 billion years old.
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 tends to be calcium-rich plagioclase
rather than alkali feldspar, the most common 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 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 unsaturated, 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 Spring Creek Shear Zone neatly divides rocks of the Yavapai
Province to the north, which are assigned to the Moppin Suite,
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 suite, such as the Moppin Suite, 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 Suite have been so heavily folded that
the older beds now lie atop the younger beds. Do not confuse a suite,
meaning a highly deformed rock unit like the Moppin Suite, with an
igneous suite, like the calc-alkaline suite, which is a
group of rocks showing distinctive chemical trends.
With that digression on stratigraphy out of the way, let's return to our story.
About 1.7 billion years ago, the juvenile crust along the
southern margin of Laurentia began to be stretched to form a back-arc
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 the accretion of 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 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. 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
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.
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.
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
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
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:
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 tetrahedon 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 ions.
Muscovite structure. U.S. Geological Survey
Muscovite is composed of triple sheets. The upper and lower layer of each sheet is a phyllosilicate layer. The layers are oriented so that they face each other, with the apical oxygen ions on the inside. Between the phyllosilicate layers is a layer of aluminum ions. The apical oxygen ions bind to the aluminum ions, which balance their charge. It's almost like a sandwich, with the two phyllosilicate layers as the bread and the aluminum as the sticky layer of peanut butter or marmite that holds the two slices of bread together.
The aluminum atoms actually have more than enough positive charge
to balance the apical oxygen ions. The balance is made up by
incorporating hydroxyl ions into the structure. A hydroxyl ion
consists of an oxygen ion bound to a hydrogen ion, with a net
charge of -1. The hydroxyl ions fit between the aluminum ions in
such a way that there is a hydroxyl ion in the center of each ring
of apical oxygen.
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 sheets have
corresponding cups that fit to the potassium ions and bind the
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 sheets. In other
words, mica crystals show 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 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 Quartize in the
Picuris Mountains give way to quartzite. Here's a sample.
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 separated out with a 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 were 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 its 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 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 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
Metarhyolite of the Vadito Group has a radiometric age of 1.70
billion years, significantly younger than the Moppin Suite. There
are particularly fine outcroppings in the vicinity of Ojo
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 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.
Within 100 million years of the deposition of the Vadito Group,
the back-arc basin began to close up, possibly helped along by the
collision of island arcs with southern Laurentia that became the
Mazazatl Province. 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 Suite 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
Part of the contact between the Tres Peidras Granite and the Moppin Suite is exposed in a roadcut just west of the village.
Here are some close ups of rocks from either side of the contact.
Contact between Moppin Suite (left) and Tres Piedras Granite (right). 36 39.155N 10558.697W
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.
South of the Spring Creek Shear Zone, the exposed Precambrian rocks are assigned to the Hondo Group. These are all less than 1.7 billion years old and are intepreted 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.
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.
Extensive outcroppings of Ortega Quartzite are found in the Tusas
and Sangre de Cristo Mountains.
Mountain in the Tusas Range is underlain by Ortega
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. 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. 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, 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.
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.
Ortega Quartzite with sillimanite vein. 36 15.776N 105 47.610W
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.
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.
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.
Muscovite from the Joseph Mine. 36 19.646N 106 03.324W
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.
Rinconada Formation. 36 15.738N 105 48.262W
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.
Staurolite is fairly uncommon. 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. But it’s another reason staurolite is uncommon.
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. 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.
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 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.
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
rounded peaks 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 Nacimeinto,
these include the oldest rocks in the Jemez region.
Relief map of the Jemez with San Pedro Quartz Monzonite outcroppings highlighted in red.
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
This outcrop has a small shear zone crossing it.
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
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.
Rapakivi texture. 36
01.354N 106 50.945W
The weathering of this surface helped bring out the texture Here’s a sample.
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 makes this a rapakivi
quartz monzonite; the oligoclase 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.
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 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 rather than
aluminum act to bond pairs of phyllosilicate sheets together. This
modifies the structure slightly. 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 crystal surrounded by eight oxygen 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.
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
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.
Relief map of the Jemez with Mazatzal province outcroppings highlighted in red.
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 part of the Yavapai-Mazazatl
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 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
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.
There is a beautiful and easily accessible outcropping of Precambrian quartz monzonite at the 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
Mesa, west of the tunnels. The tunnels and settlement were
named for William H. Gilman, the vice president of operations of
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.
The quartz monzonite itself is not much to look at when weathered.
But fresh surfaces reveal just how gorgeous a stone this is.
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
The hornblende in this rock is a very common example of an amphibole
Amphiboles are the most common examples of the family of silicate minerals called double-chain inosilicates. These are silicate minerals in which each silica tetrahedron is joined to two or three neighbors so that the silica backbone consists of two parallel chains of tetrahedra joined together:
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.
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
The monzonitic gneiss of Guadalupe Box is estimated to be about 1.6 billion years in age. Its age and location indicate that it probably should be assigned to the Mazatzal Province. The Yavapai-Mazazatl Transition Zone would then be represented by the stitching pluton of the San Miguel Gneiss.
The only Precambrian rocks exposed within the Jemez proper, in Cañon de San Diego
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. It 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 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
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.
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.
47.481N 106 41.208W
As with the leptite 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.
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 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,
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.
Relief map of the Jemez with Joaquin Granite outcroppings highlighted in red.
The Precambrian rocks of the Mazatzal and Yavapai Provinces are intruded in many locations by huge granite or granite-like bodies, called batholiths, with radiometric ages around 1.4 billion years. 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.
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
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.
Further up the road, the entire assemblage is distorted from intense metamorphism.
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. 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
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 (juvenile
crust) to mature crust. In the process, 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.
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 Suite on Hopewell Ridge.
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.
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.
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. 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. 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.
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
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
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. Here 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 of the extensive metamorphic
formations of the Tusas and Picuris Mountains, which was caused by
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 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 NaFe9(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, 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 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.
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.
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, which is a
light-colored granite containing very little dark mafic minerals.
Some quite sizable leucogranite outcrops occur in the Sierra
Nacimientos. Leucogranites are usually interpreted as a product of
the melting of thickened crust composed almost entirely of clay
Nearby is another dike of very different character.
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
Another kind of mafic dike intrudes an outcrop of the San Pedro
Quartz Monzonite along State Road 126 in the northern Sierra
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
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 risn 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 emplaced over 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.
Next chapter: When the Jemez was
Copyright ©2014-2018 Kent G. Budge. All rights reserved.