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, and
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, were a mess. They tended to be coarsely crystalline rocks, either intrusive volcanic rocks or metamorphic rocks (rocks recrystallized under great heat and pressure.) They were 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.
With the discovery of radioactivity, and 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 fully 88% of the Earth's history. With the ability to finally 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)|
We know few details of the Hadean, since no solid rocks from this
eon survive. About all we have are a few zircon grains in very old
sedimentary beds that appear to have eroded out of rocks older
than 4 billion years. During the Hadean, much of the Earth was
still molten, and there was only the beginnings of continents.
There were frequent impacts by large bodies left over from the
birth of the Solar System, including a catastrophic impact that
was the origin of the moon and a Late Heavy Bombardment around 4
billion years ago.
Archean rocks form most of the central cores of the continents, which are known as the continental shields. There is much that is still not known about this eon, including whether large continents had assembled. However, life was already present, though only as bacteria and archaea. These are two of the three domains of life found on Earth, and both have cells that lack true nuclei. The eukaryotes, whose cells have true nuclei and a complex internal structure, and whose modern descendants include animals, plants, and fungi, were probably not yet present even as simple protozoa. Oxygen was virtually absent from the atmosphere, and the Sun shone at only about 75% of its current brightness, but abundant atmospheric methane is thought to have trapped enough heat to permit the oceans to remain unfrozen.
Fossilized Precambrian stromatolites from Glacier National Park. National Park Service
The Archean bacteria included cyanobacteria, formerly known as
blue-green algae, which are capable of producing oxygen by
photosynthesis. However, the oxygen generated by Archean
cyanobacteria was removed from the environment as fast as it was
generated, by chemical reactions with reduced iron and sulfur
dissolved in the ocean water.
The earliest record of life on Earth may be deposits of graphite in the Isua region of Greenland that 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 rare 13C isotope than the much more common 12C isotope, and, as a result, biogenic carbon is depleted in 13C. 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.
The earliest generally accepted evidence of life dates to some
300 million years later. Some species of cyanobacteria form
massive colonies called stromatolites, and fossil 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 algae making up the colonies. Today, stromatolites are
found only in unusually harsh marine environments in which
predators cannot survive.
No Archean rocks are found in New Mexico, because New Mexico didn't exist yet. The oldest rocks found in New Mexico are Proterozoic rocks about 1.8 billion years old.
The Proterozoic Eon includes 42% of geologic time, and
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. The
Proterozoic also marked the emergence of eukaryotes.
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.
During the Proterozoic Eon, life began 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 kinds of 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 (amorphous 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 other signature of free oxygen is the presence of sedimentary red beds, which derive their bright red color from hematite. New Mexico, as part of the North American continent, was probably assembled at about this time, and banded iron formation is found in New Mexico in the Tusas Mountains.
We saw in the last chapter that most of the continental crust had formed by 2.5 billion years ago, which marks the start of the Proterozoic. Much uncertainty remains about the process by which this took place. It is widely believed that the continents started as small bits of continental crust that gradually assembled, sticking to each other when they were pushed together by the oceanic conveyor. But 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. While geologists disagree over whether such 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.
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 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, forming the ancient core of North America, which geologists call Laurentia. Shortly afterwards, two more provinces merged with Laurentia to form the future Wyoming and Montana region. By 1.8 billion years ago, the southern margin of Laurentia ran roughly along what is now the Wyoming-Colorado border.
During the next 80 million years or so -- longer than the time interval between today and the age of the dinosaurs -- 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 Laurentia.
A microcontinent is a small patch of continental lithosphere. The largest current 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 unusually far from the mid-ocean ridge, suggests that
this is a phenomenon of unusually cold oceanic crust. Such island
arcs consist of rock that is less dense than oceanic crust, but
more dense than continental crust.
When microcontinents or island arcs are carried into a destructive margin by the motion of the underlying oceanic lithosphere, they are unable to subduct because of their low density. If the microcontinent is quite small, the lighter crust shears off the underlying upper mantle and sticks to the continent on the other side of the subduction zone. 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 1.7 billion years ago formed the Yavapai Province, which reaches from modern-day southwest Arizona to Michigan and includes most of Colorado and the northwestern part of New Mexico. The oldest Precambrian rocks in northern New Mexico are about 1.77 billion years old.
Precambrian rocks of the Yavapai Province are exposed in the Tusas Mountains north of the Jemez. These rocks have been distorted and altered by geologic processes over the last 1.77 billion years, but geologists believe that the rocks originally consisted of thick beds of sandstone deposited on top of volcanic rocks. These subsequently experienced severe 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 minerals brought in by 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. We learned in the last chapter that mountain ranges must be held up by low-density roots, just as the tip of an iceberg is held above the surface of the ocean by the buoyancy of the rest of the iceberg lying beneath the ocean's surface. 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 show the direction in which stresses were applied to the rock while it was undergoing metamorphosis, with the foliation lying perpendicular to the direction of compression.
The mineral content of a metamorphic rock gives clues to the
temperature and pressure at which the rock underwent
metamorphosis. This is because different minerals are stable under
different conditions. Geologists speak of characteristic
combinations of minerals that point to particular temperature and
pressure regimes as metamorphic facies. I won't go into
these in any detail, because metamorphic rocks are uncommon in the
Jemez. The Jemez is mostly composed of relatively young rocks that
have not experienced deep burial.
The sandstone deposited in northern New Mexico 1.77 billion years ago 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.
Quartzite is found in great quantities in the Tusas and Sangre de Cristo Mountains, and it is believed to be part of the same formation in both locations. Geologists have named the portion in the Tusas the Ortega Formation, but it is often referred to informally as the Ortega Quartzite.
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 1.2 and 1.6 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
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 informally described using that type, as with the
Ortega Quartzite or the Bandelier Tuff.
Mountain in the Tusas Range is underlain by Ortega
Quartzite. The surrounding country exposes a variety of
metamorphic rocks of the Yavapai Province.
The Ortega Quartzite includes beds of 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 a conglomerate. An outcropping of metaconglomerate of the Ortega Quartzite can be found near the forest road on the north rim of Spring Canyon.
Ortega Quartzite also crops out in the southern Tusas Mountains. Here is one outcropping of particularly clean muscovitic quartzite.
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.
Muscovite quartzite. 36 25.682N 106 0.735W
North of Spring Canyon is Hopewell Ridge, which is composed
mostly of Moppin Series metavolcanics intruded by Marquinita
Granodiorite. These rocks originally lay beneath the Ortega
Quartzite, but the flat beds were folded up by the intense
pressure that accompanied metamorphosis, and the beds are no
longer level. A particularly interesting feature of Hopewell Ridge
is the presence of significant deposits 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 a type of metamorphosed banded iron formation, and may have been deposited as a result of the activity of cyanobacteria. The young Earth's oceans were full of ferrous iron, which was oxidized to insoluble ferric iron by the oxygen produced by photosynthesizing cyanobacteria and precipitated onto the ocean floor. It took hundreds of millions of years for the ferrous iron to be fully oxidized so that oxygen could begin accumulating in the atmosphere. The deposits of ferric iron in what is now the Tusas Mountains, which were interbedded with sediments, were later subject to metamorphism and converted to magnetite schist.
The presence of banded iron formation on Hopewell Ridge suggests
that the Moppin Series rocks were erupted in a marine environment,
as part of an island arc or in a back-arc basin.
A back-arc basin sometimes 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, 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. The Moppin Series would then be the volcanic rocks erupted in the back-arc basin as it formed, and the Ortega Quartzite would be sand deposited in the basin.
Further evidence for the formation of back-arc basins at this
time is provided by beds of pyrite-bearing chert in near
Wheeler Peak in the Sangre de Cristo Mountains. These are
thought to have formed in hydrothermal systems along the axis of a
back-arc basin. There are also well-preserved pillow basalts,
erupted under water, in the Brazos
The Moppin Series 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
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 in some detail.
We learned in the last chapter that silicon atoms prefer to covalently bond with four oxygen atoms. Each of these oxygen atoms shares a pair of electrons with the silicon, 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 need for 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 sides with opposite sides 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 boundaries 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. Note also that there is a single silicon atom on each face, with 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. Note that the silicon atom appearing as a small black sphere at the center of each unit cell is actually two silicon atoms on the top and bottom faces 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. However, the strong three-dimensional bonding gives quarts 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 ion fits snugly into one of the
openings in the structure, near the aluminum ion 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.
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 can 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 most chemical reactions, it takes placed much more quickly 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 broken thick 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 vacancy 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) are 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 it is 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 atoms, 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 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. We'll have more to say about both mica and amphiboles later in the book.
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
metashale. We'll have more to say about shales later in the book.
Leptite is found throughout the southern Tusas Mountains, though
most is not as finely laminated as our first example. In most
locations it is mapped by geologists as metarhyolite of the Vadito
Group. Its radiometric date is 1.70 billion years old,
significantly younger than the Moppin Series. There are
particularly fine outcroppings in the vicinity of the small resort
The ridge from which this panorama was photographed is northwest
of the resort, and the panorama begins to the northeast. The ridge
itself, which continues to the south in the fourth frame, is
nearly the southernmost Precambrian exposure of the Tusas
Mountains. The peak in the seventh frame 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.
To the east of the Tusas Mountains are the Picuris
Mountains, a part of the Sangre de Cristo Range. Here the
highly metamorphosed basement is exposed again. The rocks here
include include the Vadito Group, whose formations are mostly
volcanic or volcaniclastic, since heavily metamorphosed. A
volcaniclastic is a rock on the boundary between sedimentary and
igneous rock. It is formed from broken pieces of rock (clasts)
that are associated with volcanic activity.
The volcaniclastics include some beautiful metaconglomerates,
similar in age to the Moppin Series and the Ortega Quartzite.
(Some geologists regard the Ortega Quartzite and the Vadito Group
to be the same unit, exposed in different locations.)
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 Formation, 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 rock often breaks through the
clasts rather than around them, and 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
Earlier we learned that quartz is composed of silica tetrahedra, in which each silicon atom is surrounded by four oxygen atoms, and each oxygen atom joins two silicon atoms. Thus, each silicon tetrahedron is joined to four neighbors with whom it shares an oxygen atom. This forms a very strong three-dimensional network. Feldspar has the same basic structure, but with one or two out of every four silicon atoms replaced by aluminum atoms, and with extra potassium, sodium, or calcium atoms added to the crystal to supply additional electrons (since an aluminum atom has one less than a silicon atom.) All silicate minerals built on this basic three-dimensional network of interlocked silica and alumina tetrahedra are called tektosilicates.
Muscovite belong to a 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
atoms in each tetrahedra are shared; the fourth sits by itself at
the tip of each tetrahedron, as shown here. This fourth oxygen
atom is described as an apical oxygen atom. The overall structure
is of interlinked rings of silica tetrahedra.
From a chemical standpoint, this structure is incomplete. The
apical oxygen atoms are only connected to one silicon atom, and
this leaves them deficient an electron. In addition, in muscovite,
one silica tetrahedon in four is replaced by an aluminum
tetrahedron, which makes the structure even more deficient in
electrons. As with feldspar, the missing electrons are supplied by
Muscovite structure. U.S. Geological Survey
Muscovite is composed of triple layers. The upper and lower part of each layer is a phyllosilicate sheet. The sheets are oriented so that they face each other, with the apical oxygen atoms on the inside. Between the sheets is a layer of aluminum atoms. The apical oxygen atoms pick up their missing electrons from the aluminum atoms, to which they then strongly bind. It's almost like a sandwich, with the two phyllosilicate sheets 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 have more electrons to donate than the apical
oxygen atoms can accept. The balance is made up by incorporating
hydroxyl groups into the structure. A hydroxyl group consists of
an oxygen atom bound to a hydrogen atom, and it is deficient one
electron, which is supplied by the aluminum. The hydroxyl groups
fit between the aluminum atoms in such a way that there is a
hydroxyl group in the center of each ring of apical oxygen.
Each ring in the phyllosilicate sheets forms a kind of cup in the outer surfaces of the triple layer, which is lined with oxygen and hydroxyl. This forms an inviting location for a potassium atom to sit. The neighboring triple layers have corresponding cups that fit to the potassium atoms and bind the sheets together. The potassium atoms supply the electrons needed to compensate for the one silicon atom out of every four that has been replaced with an aluminum atom.
The family of phyllosilicate minerals which share the three-layer
structure of muscovite are known as micas. Different mica
minerals substitute different metals for aluminum and potassium,
but have the same basic structure.
The binding by potassium ions is not particularly strong. As a
result, mica is easily split between triple layers. In the
language of mineralogy, 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
electronics, and even as a substitute for glass.
Muscovite is an aluminum-rich mineral, with equal numbers of
aluminum and silicon atoms in its structure. This contrasts with
alkali feldspar, which has three silicon atoms for every aluminum
atom. The presence of muscovite in granite is a indication that
the granite is peraluminous, rich in aluminum.
Near the mica-rich metaconglomerates of the Picuris Mountains are
large outcroppings of quartzite, once mapped as the Marquenas
Quartzite of the Vadito Group but now regarded as identical with
the Ortega Quartzite. Here's a sample.
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. With a little coaxing, the iron atom 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. 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 atom
match the two electrons needed by each oxygen atom. Ferric oxide
has the composition Fe2O3, reflecting the
additional oxygen needed to accept the third electron from each
iron atom. 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'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
Another clue to the history of northern New Mexico in the
Precambrian is the presence of calc-alkaline igneous rocks that
are similar in age to the Moppin Series. 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, relative to its calcium content. Rocks that are enriched in the alkali metals 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. We'll have more to say about silica saturation in a later chapter.
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 the 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 west is
the valley of the Rio Guadelupe and Rio de las Vacas.
The Sierra Nacimiento is part of the Jemez region, but it has a very different origin and history. 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.
The northern Sierra Nacimiento have outcrops of a variety of Precambrian formations, which include the oldest rocks in the Jemez area. One prominent formation is the San Pedro Quartz Monzonite.
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
The outcrop has a 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. Such shear zones are
associated with deep crustal movement, with the rock on opposite
sides moving past each other. This resembles a fault, but occurs
at such great depth that the rock flows rather than fractures. In
other words, this occurs at depths below the brittle-ductile
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.
Rapakivi texture is fairly rare in granite-like rocks, but 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.
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 atoms
joining each pair of rings of the two phyllosilicate sheets, three
iron or magnesium atoms join each pair of rings. A mica having two
metal atoms joining each pair of rings is described as
dioctahedral, while one having three metal atoms join each pair of
rings is described as trioctahedral.
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.
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 new enough to be very sexy, geologists identified the Snake River Plain as a hot spot trace, where 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. However, as the rocks along the lineament were dated from radioisotopes, 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 is unusually ductile, or
hot, or contains minerals with an unusually low melting point, or
in some other way is 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 formed in the last thirty million years
and is a region of the crust stretching from central Colorado into
northern Mexico, roughly along the valley of the Rio Grande, where
the crust is slowly being pulled apart. 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. Formations 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 1.70 billion years, and there are subtle differences in isotope ratios. 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, which tend to be uncommon in oil country. 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 in great detail. 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 only Precambrian rocks exposed within the Jemez proper, in San
Diego Canyon at Soda
Dam, likely belong to the Mazatzal Province. 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 slate 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.
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 west of the Jemez. The rock here doesn't fully match the description of any of the granites further west, though its age is similar to the older granites.
Relief map of the Jemez with Mazatzal province outcroppings highlighted in red.
The Precambrian rocks of the southern Sierra Nacimiento Mountains include some that are the right age and character to be assigned to the Mazatzal Province basement. 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 carved out the rock to make way for a logging railroad. 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 somewhat
suggestive of a rapakivi texture like that of the San Pedro Quartz
Monzonite. However, the crystals are not as regular, do not show
an oligoclase rim, and are less striking. The quartz is also less
striking and this rock is more abundant in iron-rich minerals. It
has also been metamorphosed, though the foliation from
metamorphism is barely evident here.
The hornblende in this rock is a very common type of 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 atoms are required to
supply additional electrons, and, as with mica, these are
accompanied by hydroxyl groups that prevent the structure from
having too many electrons. Pairs of double chains face
each other, with the apical oxygens on the inside bonded to a
strip of metal atoms. Each such strip looks a little like an
I-beam in cross-section. The "I-beams" then interlock, with
additional metal atoms holding the "I-beams" in place.
The metal atoms holding the pairs of double chains together are
shown in gray, with the associated hydroxyls in green. The metal
atoms 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
Hornblende is a rather general term for amphiboles rich in
iron. A typical formula would be Ca2Fe5Si8O22(OH)2.
The grey atoms in the previous diagram would then be iron and the
yellow atoms 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
Like the gneiss at Soda Dam, the monzonitic gneiss of Guadalupe Box is estimated to be about 1.6 billion years in age. This is the time when the Yavapai province was being sutured to the Mazatzal province. We may be looking at a magma chamber from a volcano chain that once resembled the Cascade Mountains of Oregon and Washington.
The central Sierra Nacimiento is dominated by the San Miguel
Gneiss, which has a radiometric age of 1.695 billion years. This
puts this rock right on the geographical and chronological
boundaries between the Mazatzal and Yavapai Provinces.
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.
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 outcropping of rock that is different
in character from the rock forming the intrusion. They are also rootless;
that is, they do not connect with any deeper structure or
correlate with any nearby outcroppings. There is a single roof
pendant in the San Miguel Gneiss that caught the attention of
geologists mapping the area. It is just a couple of hundred feet
across, is located on a heavily forested hillside, and is not that
well exposed. I have to admire the geologists who were thorough
enough in their mapping to discover it.
Hornblendite outrcop in the San Miguel Gneiss.. 35 50.530N 106 51.110W
This rock is mostly hornblende, with just a scattering of feldspar. This means a low silica content, perhaps as low as 45%. Such rocks are described as ultramafic and they are fairly uncommon. Its location on the boundary of two ancient crust provinces suggests this may be a remnant of oceanic crust trapped between the Yavapai and Mazatzal Provinces when they merged 1.6 billion years ago.
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 to the south (Indian Ocean) and east (Pacific
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 old.
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 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.
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 Succession 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.
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.
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. 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 Precambrian 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 clear in this photograph, but the sample at top is nearly two inches thick.
Pegmatites that intrude amphibolite often form almandine garnet, and individual garnet crystals of some size 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 supply the necessary electrons to stabilize the structure. 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, and so we will have little more to say about it in this book, since metamorphic rock is uncommon in the Jemez.
Tourmaline is a cyclosilicate mineral, whose basic framework is stacked isolated rings of silica tetrahedra. These are joined together by triangular borate ions, while charge balance is provided by various metal ions. The type of tourmaline found at Joseph Mine is schorl, in which the metal ions are predominantly iron and sodium. The sodium and borate come from the pegmatite, while the iron comes from the amphibolite, and silica comes from both. You can see that the schorl takes the form of long black striated rods, which are particularly evident in the sample at lower right.
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. You can
see that it has a rough rhombohedral shape, and 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
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 rich in clay minerals.
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.
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 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 underplated the crust, rising to the base of the crust
because it was less dense than the upper mantle, then spreading
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, but the process is still not well understood.
The supercontinent of Columbia began to break up shortly after accreting the Yavapai and Mazatzal Provinces. This breakup had relatively little effect on New Mexico, unless the heating event at 1.4 billion years was related in some way. However, A-type granite does not have the composition geologists expect for magma associated with continental rifting.
A new supercontinent, Rodinia, began assembling around 1.1
billion years ago, and mountain building associated with
continental collisions took place in west Texas. The most likely
consequence for the Jemez area was to raise the area above sea
level and erode away any rocks younger than 1.4 billion years.
There are a couple of proposed reconstructions of the geography
of Rodinia. Both place Australia and East Antarctica along the
west coast of Laurentia, but one model has Australia far to the
north and East Antarctica directly west of the Jemez while the
other has Australia directly west of the Jemez and Antarctica to
The following animation illustrates the formation of the crust under New Mexico. It 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 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.
The timing of the breakup of Rodinia can be estimated from paleomagnetism. When magma solidifies, any magnetic minerals in the rock are aligned with the Earth's magnetic field. Unless the rock is subsequently heated above a certain temperature, its Curie temperature, this remanent magnetization remains unchanged for up to hundreds of millions of years. Geologists can measure the remanent magnetization of a rock sample and estimate the strength and orientation of the Earth's magnetic field at the time of its formation. Such paleomagnetic measurements show where the rock was relative to the Earth's magnetic poles.
Paleomagnetic measurements in Australia and North America are consistent with the continents lying side by side up to about 770 million years ago. At this time, the locations of the continents begin to diverge, suggesting that Australia had broken away from North American to produce the paleo-Pacific, or Panthalassic, Ocean.
The breakup of Rodinia around 750 million years ago, and the possible reassembly of the continents into the supercontinent of Pannotia between 600 and 550 million years ago, left no traces in northern New Mexico. The nearest rocks that record this long interval of the Earth's history are found in the Grand Canyon of Arizona, and date back around 1.35 billion years and form the Grand Canyon Supergroup exposed in the lower portions of the canyon. The period from 1.1 to 0.6 billion years ago is particularly poorly recorded throughout the American Southwest.
However, one geologic event during this time period did leave its
mark on the Jemez area.
Here are a couple of pieces of Precambrian basement rock collected from a spot northeast of Santa Fe, along the Hyde Park Road.
The rock on the right is the predominant rock in the road cut. It
is a fairly unremarkable granite gneiss, though attractive enough
that it's used for decorative gravel throughout the region. It
resembles the gneiss from Soda Dam. The rock on the left is from
the biotite schist that was intruded by the granite gneiss.
There's more to these rocks than meets the eye. The fractured surfaces, which my photos cannot do justice to, are actually quite ancient. These distinctive fractures, which are even more striking in the original outcropping, are called shatter cones, and they are an indication that the rocks were very close to a meteorite impact event. This is confirmed by findings of shocked quartz in the area. This Santa Fe impact structure is of moderate size, not more than a few kilometers in diameter, and took place long enough ago that the crater has long since been eroded away, leaving only the deep rocks that were under the center of the crater.
It is possible that some of the shattered rock produced by the
impact was preserved. About three kilometers (two miles) to the
west, at Nun's
Corner, layers of dolomite of the Mississippian Arroyo
Peñasco Group sit on top of Precambrian granite. The granite is
highly fractured, almost rotten in texture, suggesting it is breccia
(shattered rock) produced by the impact.
How long ago did this impact occur? The Precambrian granite here is between 1.6 and 1.7 billion years old. Geologists have found no trace of shatter cones in the overlying younger rocks, which are about 350 million years old. So the impact occurred sometime between 1.6 billion and 350 million years ago. In other words, we can't constrain it very well.
This was probably not a major impact event. The area over which
shatter cones and breccia are found is not large, though tectonic
activity over the millions of years since the impact have probably
erased many of the traces. But it's the only impact event
discovered so far in the Jemez area.
Next chapter: When the Jemez was
Copyright ©2014 Kent G. Budge. All rights reserved.