See also Colorado Geology Overview, The Earth At Work
Last modified 10/22/04
Skip to Igneous Rocks
When I first thought to name this article "Colorado Rocks", my hope was to limit its scope. But in truth, it's hard to name a rock Colorado doesn't have. Some of the higher-grade metamorphic rocks, like the 1.7 Ga arc-related metavolcanics, may bear little resemblance to their original state, but in one form or another, they're all here.
Petrology, the study of rocks, may sound absurd to some, but it's more practical than it sounds. Rocks hold the only available record of the history of our one and only planet. They also hold the key to two of the great pillars of human economy—mineral wealth and agricultural vigor. They tend to hold up the planet's most inspiring scenery, and they ultimately support everything we build—including houses, schools, skyscrapers, roads, bridges, tunnels and nuclear power plants.
Our biological history is also more entwined with rocks than you might think. Cell biologists studying the origins of life now have good evidence that the precursors of modern cells used rock surfaces as both cell membranes and as as catalysts for the organic reactions they required. The chimney-like mid-ocean ridge hydrothermal vents known as ^black smokers carry on such such rock-cell partnerships to this day. It's no accident that many important human enzymes and physiologically active proteins require metal ions as ^co-factors—e.g., iron in hemoglobin, magnesium in chlorophyll, chromium in insulin, copper in cytochrome c oxidase, and zinc in angiotensin converting enzyme, to name just a few.
Since mantle and lower crust rocks are only rarely exposed, the rocks of the upper crust are the main focus of petrology, even though they constitute well under 1% of the earth by volume. Of course, geologists are eager to study any rock they can get their hands on, so they particularly prize the occasional plums thrust up to the surface from lower levels.
Since the ^chemical elements are the fundamental building blocks of all ordinary materials, including minerals, let's start there. Elemental abundances at and near the surface of the earth tightly constrain the range of minerals and rocks observed, not to mention the range of possible biologic processes.
The table below lists the 14 most abundant elements in the earth's crust in decreasing order.
With oxygen and silicon alone accounting for ~74% of the crust and aluminum for another ~8%, it's little wonder minerals composed primarily of these three elements dominate the crust. They do so primarily in the form of silicate and aluminosilicate minerals built on strong chains and sheets of tetrahedral and octahedral arrays of Si-O and Al-OH bonds. In nearly all common silicate minerals, including feldspars, micas and clays, positive ions of calcium, sodium, potassium and magnesium serve both to balance out the negatively charged silicate backbones and to bind them together neatly without disturbing their crystalline structures.
It takes seven of the eight major crustal elements to fill these vital chemical roles in the sialic (Si- and Al-rich) rocks typical of the upper continental crust. Iron, the odd atom out in the major element group, figures more prominently in the higher-density minerals inhabiting the lower continental crust, the oceanic crust and the mantle. Olivine, pyroxene, hornblende and biotite are among the most common of these ferromagnesian or mafic (Ma- and Fe-rich) minerals.
Elements beyond the top 14 listed above fall into the trace element category. They're scarce in the crust for a very simple reason: By virtue of their size or charge distribution (see below), they fit poorly into the crystalline structures of typical of most crustal (silicate) minerals. The trace elements include gold, silver, copper, nickel, zinc, lead, lithium, beryllium, niobium, tantalum, tin, uranium, thorium, tungsten, zirconium and the rare earths. Many are more abundant in the mantle than in the crust. Mantle-derived, water-rich magmas are their primary means of transport to extractable crustal depths.
In 2003, geochemist Bruce Railsback published his revolutionary and ingeniously reorganized ^Earth scientist's periodic table of the elements and their ions showing not only the neutral elements but also their naturally-occurring ions. Since ions are with rare exception the stuff of earth materials, much can be learned from their habits and proclivities. Indeed, since oxygen is by far the most abundant element in both the mantle and crust, the way various cations (positively charged ions) interact with ionic oxygen constrains a great deal of geochemistry and to some extent biochemistry as well. Silicon, the 2nd most abundant element in the crust, also plays a defining role in geochemistry.
One of the most important innovations in Railsback's periodic table is the addition of contour lines of equal ionic potential—the ratio z/r of ionic charge to radius. The higher the ionic potential, the more compact or intense the ionic electric field, and the more strongly the ion interacts with nearby charge centers. Since trends in ionic abundance, mineral formation, oxide melting point, solubility, and even nutrient value all tend to follow contours of ionic potential, the new table shows at glance important chemical relationships that the standard table obscures.
Cations (positively charged ions) of low ionic potential (z/r < 4) like Na+, K+ and Ca2+ bond relatively weakly to O-2, do not form stable oxide minerals, remain in fluid phases until late in melt evolution, are highly concentrated and soluble in natural waters and serve as essential nutrients to both plants and animals.
Cations of intermediate ionic potential (z/r = 3-10) like Al3+, Fe3+ and Ti4+ bond strongly to O-2. Their compact and largely shielded charge distributions allow them to coordinate with a single negative charge center in large numbers with little mutual repulsion. Such cations tend to form stable oxide minerals, to bond in igneous minerals early in melt evolution, to concentrate in soil, to linger in the mantle, to have low concentrations and solubilities in natural waters, to collect in ferromanganese nodules on the ocean floor, and to serve inconsequential roles as nutrients.
Cations of high ionic potential (z/r > 8) like P+5, N+5 and S+6 also bond tightly to O-2 to form highly stable and soluble radicals like PO4-3, NO3- and SO4-2, but they can't form stable oxide minerals due to mutual repulsion. However, the small C+4 cation (z/r ~ 27) forms the stable oxide and greenhouse gas CO2 as well as stable carbonate oxysalts of the soluble radical CO3-2. The C+4 cation thus plays a very special role in the planet's surface temperature-regulating carbonate cycle. High potential cations share many properties with the cations of low ionic potential. Because they both readily leach out of soils due to high solubility, K+ (low potential) and NO3- (high potential) are both key ingredients in fertilizers.
The most common silicon ion, Si4+, occupies another special niche as a highly abundant cation at the cusp (z/r = 8) between high and intermediate ionic potentials. Thanks to similar ionic potentials, Si4+, V5+, Mo6+ and Se4+ all stand at the upper margin of cations forming stable soluble oxysalts—e.g., silicate, SiO4-4 or Si(OH)4—that also form stable insoluble oxide minerals—e.g., silica, SiO2, as in quartz. (Interestingly, these 4 cations all serve as essential vertebrate micronutrients.) However, the crustal abundance of Si4+ far exceeds that of all the others in this group. Thus Si4+ appears in large quantity in both the insoluble products of weathering, most notably as sand, and in natural waters as dissolved silica. Si4+ binds to igneous minerals only at intermediate to low temperatures and remains abundant in fluid phases to the end of the crystallization sequence. Along with their low densities, these properties insure the crustal accumulation of quartz and silicates during the earth's chemical differentiation.
Another important innovation in Railsback's revamped periodic table is the grouping of ions according to their electronic configurations as ions, a dimension separate from ionic potential. Having lost all their outer shell electrons, the hard cations are left with a relatively inert noble-gas-like electronic configuration, while the soft cations retain some outer shell electrons—the more, the softer. Hard and soft cations behave quite differently. Hard cations like Ca2+ coordinate strongly with O-2 and F-; soft cations do not. When they form oxide minerals, hard and soft cation oxides have high and low melting points, respectively. Soft cations like Cu+ bond with S-2 and the larger halides I- and Br- rather than with O-2 and F-; they tend not to form oxide minerals. Thus hard Ca2+ forms both oxide and an oxygen-rich sulfate (gypsum, CaSO4) but not a sulfide, while soft Cu+ forms a sulfide (chalcocite, Cu2S) but not an oxide.
The metal cations commonly found in silicate minerals (Na+, K+, Ca2+, Mg2+) are all hard. Their low ionic potentials and noble-gas-like electronic configurations allow them to fit cleanly between large polymeric silicate and aluminosilicate sheets and chains and bind them together without disturbing them. The trace elements, on the other hand, generally have low to intermediate ionic potentials and soft to very soft outer shell configurations. Due primarily to the latter, they fit poorly in silicate lattices and for the most part remain sequestered in the mantle, where more hospitable non-silicate minerals dominate.
The most commonly occurring anions (negatively charged ions) are O-2, S-2, Cl-, F-, and the soluble oxo complex radicals CO3-2, SO4-2, SiO4-4, NO3- and PO4-3. Hard cations prefer to coordinate with O-2, by far the most common anion in both crust and mantle, while soft cations like Pb2+, Cu2+ , Zn2+ and Ag+ instead prefer S-2 (number 13 of the 14 most common elements. This preference alone accounts for some of their rarity. The extremely soft Au+ cation can't form an oxide and can only form a sulfide with the help of other soft cations—hence the long-admired rarity and "nobility" of gold and its predilection to go native in elemental form. Not surprisingly, metal oxides and sulfides are the most common ore minerals in the Colorado Mineral Belt and elsewhere. After O-2 and S-2, the properties of the anions appear to be less important than those of the cations.
Now that we've seen how crustal elements combine to form minerals, let's look at the rocks the minerals make.
Every grade-schooler knows that rocks come in three basic flavors—igneous, sedimentary and metamorphic, as detailed in the table below. That's still an excellent starting point, but we'll need some subtypes to make real headway in understanding the rocks of Colorado. It's worth emphasizing at the outset, however, that rocks in the field form a continuum of origins, compositions and textures beyond the reach of any rigid classification scheme. No matter how fancy the classification, there will always be important transitional rocks that can and will be classified more than one way by reasonable geologists.
Much to the dismay of architects, students and users of rock classifications, transitional rock types pop up everywhere. Important examples include the following:
Confused? Hang around with rocks long enough, and you'll get used to it.
Rocks composed entirely of interlocking crystals of one or more minerals are said to be crystalline. All unequivocally igneous rocks are crystalline, as are higher-grade, recrystallized metamorphic rocks like gneiss and schist. Strictly speaking, some chemical sediments like limestone and chert (microcrystalline quartz) are also composed of crystals, but in common usage, the term excludes sedimentary rocks.
Mechanically, crystalline rocks tend to be stronger and more resistant than other types. They form the crests of the Rockies' highest ranges and hold up most of Colorado's Fourteener summits.
For the most part, rocks are equilibrium products relatively stable at the conditions under which they formed but chemically or mechanically unstable in all other environments. Crystalline rocks formed at depth are unstable at the surface, while sedimentary rocks formed at the surface are unstable at depth.
Rocks thrust into new settings by tectonic, magmatic or erosional events tend to move around the rock cycle (below) from one basic rock type to another. For example, igneous rocks formed at depth under high temperature (T) and pressure (P) in the absence of free oxygen and water are bound to change when brought to the surface to face chemical and mechanical weathering, erosion, transport, deposition and diagenesis. Given enough time, their debris will become sedimentary rocks best suited to surface conditions. With deep burial during a mountain-building event, the elements in the sedimentary rocks will reorganize into new metamorphic minerals better suited to the extreme pressure-temperature (PT) conditions they now face. Under the right PT conditions, they might even come full circle to melt back into igneous rocks. Alternatively, uplift and erosion of the metamorphic rocks might ultimately produce a new batch of sedimentary rocks. All paths through the rock cycle are possible.
The diagram at right nicely summarizes the rock adjustments outlined above. It's useful to think of igneous rock as the starting point for the cycle, but material can jump in anywhere and end up anywhere. The transformations that occur with any frequency are already shown in the diagram, but since deeply buried sedimentary rocks can melt directly in certain tectonic environments, it would be reasonable to add another thin black curved arrow pointing from the sedimentary to the igneous node.
Acknowledgment: Rock cycle diagram courtesy ^Lynn Fichter.
Rocks that are slow to weather and erode are said to be resistant. All other things being equal, erosion will leave resistant rocks standing higher than the less resistant rocks around them. Crystalline (igneous and metamorphic) rocks then to be more resistant than unaltered sedimentary rocks, but chert is a notable exception.
Fusibles are rocks or minerals that melt easily. Refractories are just the opposite. Sedimentary rocks tend to be fusible, while crystalline rocks tend to be refractory, some more than others. Not surprisingly, the most refractory rocks, like gabbro and peridotite, reside in the lower crust and mantle. Sedimentary rocks groomed for surface stability wouldn't stand a chance at those depths.
Rock materials don't necessarily move around the rock cycle as the rocks they compose evolve. In a tectonic disturbance, uplifted sedimentary rocks can be reworked into new sediments, igneous rocks can remelt, and metamorphic rocks can prograde. Reworking adds yet another layer of complexity to rock genealogy.
Many earth processes play out at depth beyond the reach of the atmosphere and hydrosphere, but for many others (including weathering, erosion, transport, deposition, isostatic rebound and basin subsidence), the rubber meets the road at the surface, where the atmosphere, the hydrosphere and the land all interact strongly to shape both land and climate in a never-ending dance.
Once weathering gets a foothold on a rock exposure, erosion, transport and deposition are likely to follow, but weathering continues to break down the sediments, both en route and at the site of deposition. Given sufficient time and transport distance, the ultimate end-products of weathering are always pretty much the same, regardless of the initial rock type:
Why end up with just these three mineral groups? Because all the rock-forming minerals commonly exposed on this silicate planet eventually break down into quartz, clay or calcite unless some other process (like melting) intervenes. These end-products are chemically stable under most surface and near-surface conditions, but their precursors are not.
Quartz (SiO2) grains are exceptionally stable at the surface. They may be ground down to silt-size during transport, but like glass (also SiO2), they're chemically inert. (That's why chemists use glass containers.) Once silt-sized, they go back into suspension in moving water, where they escape further mechanical weathering.
Clay minerals tend to form microscopic flat platy crystals with charged surfaces that slide easily against each other and have a hard time interlocking, especially when wet. Claystones tend to be weak as a result, and clay particles remain in transport the longest because of their size and shape. The most common clay minerals produced by weathering—montmorillonite, illite, kaolinite, in order of current abundance—reflect the stability of sheeted Si-O and Al-OH crystal structures. Montmorillonite is the expansile clay dreaded by homeowners and civil engineers everywhere. Kaolinite formation is restricted to low latitudes because it requires a hot wet climate.
Some sections in this article close with a "Map Units" subsection describing how to find pertinent bedrock (surface rock) units on the Geologic Highway Map of Colorado.
Rocks that solidify from a molten or partially molten state are said to be igneous. Rocks that freeze on or above the surface, whether in the air or underwater, are extrusive or volcanic. But if they freeze below the surface for any reason, as did the 1.4 Ga Silver Plume granites exposed so handsomely at ^Rocky Mountain National Park (right), they're called intrusive or plutonic instead. As we'll see, extrusive and intrusive igneous rocks differ chemically, texturally, and in other important ways. In either case, magma is the molten rock involved. Magmas reaching the surface in liquid state are called lavas.
The geothermal gradient guarantees that all melts develop at depth. Since melts are almost always lighter than the solid rocks from which they derive, gravity impels them to rise toward the surface as best they can, just as a bubble eventually rises, however slowly, through semi-solid molasses in the refrig. (A rising rock body, whether solid or molten, is a diapir.) In fact, most rock melts are buoyant enough to approach the surface if they don't freeze first.
On the way up, melts interact both physically and chemically with the rocks through which they pass. In the process, they give off heat and fluids and take in easily melted or dissolved wall rock components. The final igneous product, whether extrusive or intrusive, is always highly evolved relative to the initial melt, but some magmas reaching shallow levels are more primitive than others. On average, the basalts erupted at seafloor spreading centers are the least evolved relative to their asthenospheric source rocks. Continental granites and rhyolites are among the most differentiated of all magmas.
Since the onset of plate tectonics ~2.0 Ga, most of the planet's igneous activity has concentrated along plate boundaries. As the modern map below clearly shows, the situation is no different today. Unusually eruptive boundary segments like Iceland are called hot spots. The igneous centers found far from plate boundaries are also hot spots (Hawaii is the only one shown here, but others exist.) Hot spots arise for a variety of reasons, most of which ultimately relate to extensional failure of the plate(s) involved; the once-popular deep mantle plume explanation for hot spots is not supported by the observations. Volcanic outputs at hot spots can be truly prodigious, but the associated intrusive activity can be just as important.
In many ways, igneous rocks are the starting point for the rock cycle. Whether the parent melt derives from the mantle or from sedimentary rocks buried deep in the upper crust, it eventually cools and freezes into a solid mass of interlocking crystals derived from a handful of mineral families, including those listed in the table below. The igneous raw materials can then go on to become sedimentary or metamorphic rocks as events and conditions unfold.
* Table Note: Since quartz and feldspathoids (AKA foids) are chemically incompatible, they seldom occur in the same rock. Most of the remaining possible mineral family combinations can and do occur in common rocks.
The term felsic means feldspar- and silica-rich. Sialic rocks (those rich in silica and aluminum) are particularly felsic. Felsic rocks tend to be of continental origin. Felsic magmas like rhyolite have typically reacted strongly with continental (or at least felsic) crust on their way to the surface, regardless of the source of melt.
Mafic means Mg- and Fe-rich. Rocks of the upper continental crust are felsic on average, but mafics are fairly common there. Rocks of the lower continental crust and the oceanic crust are almost always mafic. Mafic rocks are on average denser than felsics and tend to be found at greater depths as a result. Ultramafic rocks are exceptionally rich in Mg and Fe and poor in Si. They almost always come from the mantle and accordingly tend to be very dense.
A single melt can produce a wide variety of different igneous rocks through a complex process known as magmatic differentiation. Exposed magma bodies, whether intrusive or extrusive, typically display some degree of differentiation over both space and time. Most melts develop in the lower crust or in the upper mantle's asthenosphere. As a result, many melts start out with a fairly primitive mafic or basaltic composition. Melts developing in the upper crust tend to have higher initial silica contents
Regardless of where they form, all melts evolve considerably during ascent. The rocks they ultimately produce depend on the composition of the original melt and on the properties of the wall rocks encountered en route. The main processes involved are fractional crystallization, assimiliation, exchange of volatiles, and magmatic mixing.
Fractional crystallization (or fractionation for short) occurs when circumstances prevent early-forming crystals from reacting with the remaining melt. This process accounts for most of the differentiation observed in igneous rocks.
As a rising melt cools and reacts with surrounding rock, the melt minerals with the highest melting points or the lowest solubilities (the refractories, like olivine and pyroxene) crystallize out first, while those with the lowest melting points or solubilities (the fusibles, like silica) freeze out last. Heat released by the crystallization of refractories replaces heat lost to the surrounding country rocks by simple conduction, by country rock melting and by the assimilation of country rock fusibles. Fusibles and refractories enter and leave the melt at specific temperatures and pressures, which tend to occur at specific depths along the ascent. As it continues to rise, the surviving melt loses volume, and its fusibles become more and more concentrated. It leaves behind a trail of solid refractories and country rock alterations.
Gravitative differentiation, the most common form of fractionation, stems from the fact that most solid minerals are more dense than their parent melts. When their crystals settle to the bottom of the magma body, they are effectively segregated from the residual melt. Rocks formed from crystals amassed in this manner are called cumulates, and they're often zoned, with first crystals to leave the melt at the very bottom of the magma chamber. Cumulates formed from lighter crystals that occasionally precipitate out of the melt float to the top instead, with the lightest at the very top. Cumulate crystals are typically cemented by residual magmatic fluids.
Ascending magmas also evolve chemically by recruiting easily melted or dissolved components (fusibles) from the walls of their conduits. Heat and magmatic fluids mediate the process. In so doing, they may pick up volatiles, extra silica, trace elements and even chunks of wall rock. The thermodynamics and geochemistry involved are exceedingly complex, but the heat the melt gains from leaving behind refractories (an exothermic process) is usually sufficient to cover the heat lost to the endothermic reactions involved in the assimilation of country rock components. Assimilation can thus proceed without tapping the heat required to keep the melt from freezing.
Wall rock chunks that survive more or less intact, without completely melting or dissolving into the magma, are called xenoliths. Surviving wall rock crystals are called xenocrysts. Together, xenoliths and xenocrysts provide invaluable information about the rocks residing at rarely exposed lower crust and mantle levels.
Figuring prominently in the process of assimiliation are the volatiles found in varying amounts in nearly all wall rocks and magmas—CO2, SO2, O2, Cl2 and most notably, H2O. Water is particularly available in wall rocks of the mid-crust, both in free form and within the hydrated minerals commonly found at such depths. Some of the assimilated water goes into hydration reactions with predominantly anhydrous melt components, but most of it just builds up in the ever-shrinking surviving silicate melt. If it takes on enough water, the melt will eventually develop a water-saturated silicate fraction and a separate water-based fluid phase.
Under certain conditions, the water-saturated silicate fraction can give off a whitish fine-grained vein-filling slurry of quartz and feldspar known as aplite. The water-based phase easily assimilates trace elements that don't fit well in most silicate crystals, including lithium, beryllium, niobium, tantalum, tin, uranium, thorium, tungsten, zirconium and the rare earths. Many ore deposits form when this hot, pressurized, mineral-laden hydrous fluid finally permeates fractured country rock and cools into veins of pegmatite—an igneous rock containing unusually large crystals of quartz, feldspar and, now and then, highly prized minerals as well. Pegmatite and aplite dikes and veins are common around intrusions. Pegmatite is the prospector's friend.
The mixing of two separate magmas just before eruption or final subsurface emplacement is uncommon, but in areas of active magmatism, adjacent magma bodies are bound to develop transient subsurface communications now and then. At right is a rare banded tuff from the Valley of 10,000 Smokes, Katmai, Aleutian Archipelago, Alaska. The banding reflects the last-minute mixing of lavas from two separately differentiated magma chambers underlying the valley during the cataclysmic 1912 eruption of Novarupta Volcano, which released a whopping 30 km3 of pyroclastic material at the time. [banded tuff photo]
A much more common form of magmatic mixing involves the secondary melting (anatexis) of mid- to lower crustal rocks on contact with much hotter rising mafic melts of mantle origin to produce felsic (feldspar- and quartz-rich) magmas in arc and continental rift settings. On reaching high crustal levels, such melts may arrive with more mantle heat than mantle material in tow.
When magma reaches the surface, the excess volatiles escape in vapor form. Gases usually boil out of low-viscosity basaltic lavas relatively peacefully, as they usually do in Hawaiian eruptions. A good example shown at right is the ^March 26, 1984 fissure eruption on Mauna Loa's Northeast Rift Zone. But volatiles are more likely to explode than boil out of viscous lavas like rhyolite and andesite, as they did at Mt. St. Helens on May 18, 1980 (shown at the top of the next section). Volcanic habits are discussed in greater detail below.
Igneous rocks that solidify from melt on or above the surface of the solid earth are called volcanic after Vulcan, god of fire. The term extrusive is synonymous with volcanic. At right is the explosive June 12, 1991 eruption of Mount Pinatubo, Luzon, Philippines.
Volcanic rocks occur in many tectonic settings, including magmatic arcs at subduction zones (as in the Banda Sea at right), seafloor spreading centers, ocean islands, and along continental rifts and other leaky faults. Like their intrusive counterparts, extrusive rocks are categorized primarily on the basis of texture and composition.
Because volcanic rocks tend to cool quickly after eruption, individual mineral crystals have little time to grow and usually end up too small to see with the unaided eye. The resulting rock texture is said to be aphanitic. Occasionally, one mineral, often a feldspar, manages to grow phenocrysts (large crystals much bigger than all the rest) before venting. An otherwise aphanitic volcanic rock containing phenocrysts is called a porphyry; the fine-grained component is called the groundmass. Rock textures in which two very different grain sizes predominate are termed porphyritic. Volcanic glasses like obsidian tell of ultra-fast cooling rates.
Lava flows are perhaps the simplest of volcanic deposits, but they show their share of complexities. Between the frozen gas bubbles (vesicles), if any, most lavas are predominantly aphanitic in texture, but porphyries also occur. Over time, flows tend to vary in texture and composition, in part because they tap different portions of the magma chambers that feed them. Flows often cross eroded surfaces and interact with their soil covers and groundwater along the way. Lavas quickly chilled in air or water develop glassy textures. ^Obsidian (usually of rhyolitic composition) and the glassy rinds on basaltic pillow lavas are examples.
At right, Mauna Loa looms over Kilauea Caldera at ^Hawaii Volcanoes National Park as fume rolls off Steaming Bluff in the morning light. Mauna Loa is the world's largest mountain and largest volcano. Kilauea is the world's most active volcano. Basaltic lavas built both just in the last 1 Ma. Olivine basalt featuring macroscopic green olivine porphyrocrysts in a black groundmass is a common lava around Mauna Loa. Waves and currents have concentrated olivine dense crystals weathered out of sea cliffs at the southern tip of the Big Island into a unique green sand beach.
Solids thrown from a volcanic vent are called ejecta, and accumulations of ejecta are called tephra or pyroclastic deposits. Pyroclastics come in many sizes: Blocks and bombs are over 32 mm in diameter, with bombs showing some degree of aerodynamic rounding; lapilli are 4-32 mm across; and ash particles are under 4 mm. Wind can carry fine ash hundreds of kilometers, but, not surprisingly, larger and larger ejecta fall progressively closer to the vent. Tuff, rock made from consolidated ash layers, comes in water-laid and air-fall varieties. When ash is hot enough and falls at high enough rates, individual particles can fuse on burial by subsequent ash falls. An extremely resistant welded tuff or ignimbrite results.
Around 30 Ma, welded tuffs blanketed the entire Basin and Range to great thicknesses in a prolonged and undoubtedly unpleasant event known as the Ignimbrite Flare-up. During the Early Phase of Tertiary magmatism in Colorado (~37 Ma), massive ash flows from the Mount Princeton area rolled 90 km across the Eocene erosion surface to blanket the western Denver Basin with incandescent ash which compacted and fused to become the welded Wall Mountain tuff, apparently in a single day. Given the 10:1 compaction ratios typical of welded tuffs, modern Wall Mountain remnants up to 40' thick imply initial ash deposits up to 400' thick.
I keep the clast of Wall Mountain tuff pictured at right on my desk to help me remember what a really bad day looks like.
Volcanic rocks vary widely in their elemental and mineral content. Silica content is perhaps the important single compositional property because it strongly controls viscosity. Viscous high-silica lavas like rhyolite and andesite tend to erupt explosively, because stickier lavas retain more of their volatiles until they near the vent and release them more violently on eruption. As with crustal rocks in general, silica (SiO2) and alumina (Al2O3) dominate all volcanic rocks, even the most mafic basalts, but the silica variations shown below are more than adequate to create big differences in lava viscosity.
In order of increasing silica content and explosive tendency, basalt, andesite, dacite and rhyolite are the most common lavas. As the lava of choice of at mid-ocean ridges, on seamounts and submarine plateaus, along continental rifts and other leaky faults, basalts outnumber and outmass all other lava types by a wide margin. Basalts are mafic (rich in ferromagnesian minerals), whereas rhyolites are felsic (rich in high-silica feldspars, feldspathoids and quartz). Andesites and dacites are in between. Compositionally, rhyolite is the extrusive form of granite, while basalt resembles the intrusive rock gabbro.
Table note: The external links above lead to excellent rock photographs and descriptions provided by the US Geological Survey Photo Glossary of Volcano Terms.
For the visual learner, the diagram below shows much the same information.
The way a particular volcano behaves depends on many things, including
Of these, the most important influences relate to the properties of the lava itself.
Immature oceanic arcs usually start out with basaltic lavas. Seafloor spreading centers, ocean island volcanoes, volcanic seamounts and volcanic submarine plateaus overwhelmingly produce basalts. Basalt is also the magma of choice in continental rift settings, at least for starters, and along leaky faults of all kinds.
Subaerial basaltic volcanoes generally have limited potential for devastation at a distance. Their lavas tend to erupt at high temperatures with low viscosities, the latter due to their low silica contents, and their products tend to accumulate near the vent. That makes for fairly peaceful eruptions and for relatively stable volcanic edifices, including shield volcanoes like Hawaii's Mauna Loa, shown at right in a 1984 fissure eruption. Note the calm observer in the lower left corner of the photo. Standing that close to an erupting vent would be totally insane on a non-basaltic volcano.
Still, even basaltic volcanoes have their moments. Groundwater meeting hot rock beneath the Kilauea Caldera caused a massive week-long phreatic (steam-driven) eruption that blasted Halemaumau crater out of the caldera floor in 1924. When the dust had settled, the new crater was 915 m across and 550 m deep, and 750,000 m3 of ejected debris littered the caldera floor, including many angular car-sized blocks.
Andesites, dacites and rhyolites are the magmas most commonly erupted at magmatic arcs associated with subduction zones, particularly those built on continental margins. These viscous lavas tend to build unstable stratovolcanoes—like Mt. Rainier (right) of the Southern Cascade Range—prone to release massive lahars (volcanic mudflows) and ground-searing ashflow (glowing cloud) eruptions. As Mount St. Helens so vividly reminded us, they also tend to erupt explosively, potentially spreading large volumes of ash over wide areas. The apocalyptic "X" bentonite layer of the Denver Basin records the fall of 15 m (48') of ash from an Idaho vent hundreds of kilometers upwind.
Most arcs and many rifts go through the full range of magmas—first basalt, then andesite, then dacite and finally rhyolite—as source melts interact with wall rocks on the way up to their reservoirs, and as the magma reservoirs mature over time. But volcanic centers in both rift and arc settings not uncommonly jump straight from basaltic to rhyolitic products. If intermediate magmas are largely absent, the eruptive sequence is called bimodal. Bimodal sequences used to be taken as markers for continental rifting, but many bimodal arc sequences have been recognized as well.
Clastic sediments derived from solid volcanic source rocks are called volcaniclastic and are usually classified and mapped as sedimentary rather than igneous. Volcaniclastics and their metamorphic derivatives deserve special mention here as an important transitional rock type in Colorado's physical evolution.
Clastic volcanic sediments can accumulate in large volumes around active magmatic arcs and volcanic fields. Their fine-grained feldspar and ferromagnesian minerals quickly weather to clays, and their free quartz crystals are already silt-size or smaller. Arkose (feldspar-rich) and lithic sandstones are common in immature volcaniclastic sediments deposited near the source, while silty mudstones are common among mature volcaniclastic sediments. Immature volcaniclastics shed to the east from early Laramide volcanic highlands in the Front Range uplift accumulated to form the Denver, Arapahoe and Dawson Arkose formations at the top of the Denver Basin.
Under regional metamorphism, volcaniclastics tend to evolve along the shale pathway detailed below. In metamorphic form, arc-derived 1.7 Ga volcaniclastics make up a good share of Colorado's Precambrian basement, including the craggy hornblende gneiss exposed atop Royal Mountain (right).
On the Geologic Highway Map of Colorado, the surviving volcanic outcrops are all Tertiary in age. The Oligocene andesites labeled "Tov" in brown are by far the most extensive and occur primarily in the San Juan and Thirtynine Mile volcanic fields. Widespread Miocene-Pliocene basalts and bimodal products are mapped as "Tuv" in orange. The legend shows a Quarternary basalt unit "Qv" in stippled orange, but I have yet to spot a mapped occurrence—not even at the Dotsero Volcano, last active about 4 Ka. (Tweeto's 1979, 1:500,000 Geologic Map of Colorado correctly shows a Quaternary basalt unit there.)
Skip to Sedimentary Rocks
By definition, intrusive igneous rocks solidify from a melt before reaching the surface. The term plutonic (after Pluto, god of the underworld) is synonymous with intrusive. Coherent bodies of plutonic rock are called intrusions. The host rocks surrounding the intrusion are referred to as country rocks. Intrusives are often involved in the plumbing system of a volcanic edifice or field, as were the beautifully exposed Oligocene intrusions seen at right in the West Elk Mountains, but some never communicate with the surface while still molten. Like their extrusive counterparts, intrusive rocks are categorized on the basis of texture and composition.
The highest point on the horizon in the photo at right is Mount Evans (14,264') as seen from Denver. Its summit area exposes a large intrusion (batholith) of resistant granite emplaced during the 1.4 Ga Berthoud Orogeny, perhaps the most common intrusive rock type found in Colorado.
Slower subsurface cooling times lead to larger average intrusive grain sizes—much larger than the microscopic grains typical of extrusive (volcanic) rocks. Magmas solidifying underground tend to cool very slowly, in part because the already warm solid country rocks surrounding the intruding magma are effective thermal insulators. Individual crystals commonly reach 2-4 mm diameters, but in pegmatites, the quartz and feldspar crystals can exceed 100 mm. Such igneous textures are called phaneritic.
The grains in the granodiorite bolder at right are easily visible with the unaided eye. This particular boulder came from an Oligocene intrusion exposed along the lower Cathedral Peak in the Elk Range.
Most intrusive rocks contain the minerals shown in the diagram below. From left to right in the mineral plot at the bottom of the diagram, the rocks go from mafic (Mg- and Fe-rich) to felsic (feldspar- and silica-rich) in composition.
A mix of light and dark grains is typical in intrusive rocks. The light-colored grains include quartz (usually colorless to gray) and one or more feldspars (off-white to pink, green or gray). Micas, if present, tend to be either black (biotite) or white (muscovite). Olivine, pyroxene and hornblende range from green to black. In intrusive rocks, all these minerals have characteristic crystal shapes readily observable with a hand lens when present, but the individual crystals aren't always well developed.
Intrusive rocks are easily recognized as such but difficult to classify, particularly in the field. The classification schemes used by geologists are far too complex to discuss there, but the most common intrusive rocks are worth exploring.
Most of the intrusives found in Colorado are light-colored felsic (feldspar- and silica-rich) to intermediate rocks along granite-granodiorite lines. Many of the Colorado rocks called and even mapped as "granites" turn out to be granodiorites or other felsic types on closer inspection. A few mafic (Mg- and Fe-rich) Laramide intrusives have been found in Colorado, but mafic intrusions are not uncommon worldwide.
Table Note: The external links above lead to Lynn Fichter's well-illustration igneous rock descriptions.
Depending on the composition of the magma, the final depth and the nature of the country rocks, intrusions can take many forms.
On the Geologic Highway Map of Colorado, 1.7 Ga granites are marked "Xg" and mapped in light gray. The 1.4 Ga Berthoud and 1.1 Ga Grenville granites are lumped under "Yg" in a darker gray. The Laramide intrusions are tagged "TKi" and mapped in a dark purple, while the mid-Tertiary intrusions are marked "Tmi" in hot pink. The TKi units are pretty much confined to the Colorado Mineral Belt; the Tmi units are also heavily concentrated there but have a much wider distribution, including an occurrence at Spanish Peaks. Younger intrusives under 20 Ma labeled "Tui" in stippled pink mainly occur north of Steamboat Springs.
Skip to Chemical Sedimentary Rocks
Sedimentary rocks form from
Because they're a lot more stable under surface conditions than igneous or metamorphic rocks, sedimentary rocks now cover two-thirds of the total area of the continents to an average depth of ~0.5 km (1,800'). By volume, sedimentary rocks are about two-thirds mudstones world-wide. In Colorado, the sedimentary cover ranges from absent over the Precambrian cores of Laramide uplifts to 4.0 km (13,000') thick at the west end of the Denver Basin.
It will once again prove useful to categorize sedimentary rocks, but bear in mind that there will always be transitional rock types. Nature, in her devotion to entropy, just loves to mix things up. This section will continue with the dominant class of sedimentary rocks known as clastics. The following section will describe the chemical sedimentary class—a very different kettle of rocks indeed.
For more on sedimentary rocks, visit the extensive and well-illustrated ^sedimentary rock site by educator and geologist ^Lynn Fichter. They're not as dull as you might think. For an introduction to stratigraphy, the science of reading stacks of sedimentary rocks, see Formations and Sequences below.
Clastic sedimentary rocks form from clasts (fragments) weathered or otherwise disaggregated from other rocks. A synonym for clastic is detrital. Clastic rocks can be categorized in any number of ways, but most classification schemes key on both grain size and composition, which together tell a lot about the source rock, the depositional environment, and all the steps in between. Clastic rocks are generally held together by the interlocking of grains and by chemical cements, which also add color. Sandstones, shales and carbonates like limestone account for over 95% of all sedimentary rocks because they're the most stable rocks at surface conditions. When the source rocks are volcanic, the sediments are said to be volcaniclastic.
I can't improve on the sedimentary classification web sites developed by geologist and educator ^Lynn Fichter, so I'll just link his explanation of the ^QFL (quartz, feldspar, lithic) naming system here. The QFL system takes into account both texture and composition, but its use is beyond the scope of this site.
Clastics can be subdivided and are often named according to their average or dominant particle size according to the logarithmic Wentworth Grain Size Scale summarized in the table below.
Technical Note: The ø (phi) scale facilitates grain size statistics; ø = -log2 mm. It's worth repeating that the Wentworth scale is about grain size, not composition. Any particle in the 0.0625-4.0 mm range is sand, whether it's made of pure quartz or feldspar or undisaggregated rock fragments. In the last case, we might speak of a lithic sandstone, but it's a sandstone nonetheless. The same applies to the clay category, but here, size implies composition: Few clastic sedimentary rocks composed of clay-sized grains will contain anything other than clay.
Common sedimentary rock names like sandstone, mudstone, siltstone, and claystone come directly from their average grain sizes, as defined in the Wentworth grain-size scale above. Thus, a rock with grains averaging 0.375 mm across is a medium sandstone.
Shale, a term synonymous with mudstone, includes any fine-grained clastic sedimentary rock originally made of mud. Thus, siltstones and claystones are both shales. Pelite is yet another name for mudstones. Wind-deposited silts bears the special name loess.
Of course, all size combinations are possible, including sandy siltstones and silty claystones. The term wacke lumps together silty and shaley sandstones, which can be hard to tell apart in the field. If a gently bitten shale feels gritty between the teeth, it's one-third to two-thirds quartz silt and may qualify as a wacke. If it feels smooth or creamy, it's at least two-thirds clay.
A conglomerate is dominated by rounded gravel-sized clasts (> 2 mm), while a breccia contains angular gravel. At right is the basal layer of the Fountain Formation at Red Rocks Park, a moderately coarse conglomerate containing rounded pebbles and granules up to ~50 mm across in a sandy matrix—a typical terrigenous deposit. Since such deposits develop in alluvial fans, they're sometimes called fanglomerates.
The standard deviation of grain sizes found in a clastic rock is a measure of its degree of sorting, which in turn says something about the rock's porosity, permeability and depositional history. In the field, the standard deviation is half the size range that includes two-thirds of all the grains. If two-thirds of all the grains in our medium sandstone above fall between 0.5 and 0.25 mm, the standard deviation is (0.50-0.25)/2 = 0.125 mm. Clastic rocks deposited near their source tend to have poorly sorted grain sizes. Well-sorted rocks are both more porous and more permeable than poorly sorted rocks with the same average grain size. The degree of sorting is a measure of the textural maturity of a clastic rock.
Grain shapes also tell a story. Grain shapes are usually described in terms of their sphericity (the degree to which all dimensions are equal) and roundness (the lack of sharp corners). Sphericity is often controlled by composition. Grains can be rod-shaped, disc-shaped, or spherical. The degree of rounding depends on grain size and hardness and on details of transport, deposition and diagenesis.
Sedimentary rocks typically include grains of the following components:
Most sedimentary classification schemes key on the relative proportions of quartz, feldspar, lithic fragments and matrix. The ^QFL (quartz, feldspar, lithic) naming system is such a scheme.
Ever since plants became abundant in the Silurian around 420 Ma, it's been hard to find a clastic rock devoid of organic content, but organics usually don't figure explicitly into the classification of clastic rocks, however important they may be to petroleum geologists. At right, organic carbon darkens some of the layers in the otherwise light-colored South Platte sandstone member of the mid-Cretaceous Dakota Group.
Note that grain size and composition aren't entirely independent. Quartz grains battered down to silt size go into suspension in flowing water, where they are largely immune from further diminution and rounding by impact and abrasion. At maturity, quartz grains tend to stabilize at silt size with some residual angularity.
Sedimentary structures reflect depositional heterogeneities at scales larger than texture. They provide valuable and often quite specific information about the depositional environment.
Bedding is the most obvious example of structure. Bedding planes reflect some change in the depositional environment, whether it be
The famous I-70 road cut west of Denver (above) exposes obvious bedding in Cretaceous and Jurassic sediments of the Dakota Group, Lytle Formation and Morrison Formation, which go from left to right in this picture.
Graded bedding with grain size coarsening toward the base is usually seen in turbidites—marine sediments deposited by turbidity currents in fans on or near the continental rise. As the current slows, the larger, heavier grains drop out first. The cycle repeats with each new turbidity current. The diagrams at right show the basic features of this process, which operates along passive continental margins and magmatic arcs as well.
Cross-bedding is a common structure in sandstones and siltstones, whether water- or wind-laid. Sand dunes, beaches, tidal flats and river bars all produce cross-bedding of varying steepness and scale, but individual beds usually dip down the current that deposited them. The 1.7 Ga Coal Creek quartzite still shows delicate cross-bedding after all these years.
Ripple marks are the first bedforms to appear when fluid (air or water) flow becomes fast enough to transport sand. Symmetrical ripple marks like those exposed in the Dakota sandstone on Dinosaur Ridge (right) tend to form in lakes and on beaches, where water currents are bi-directional. Asymmetric ripples formed by streams or winds are steeper downstream or on the lee side.
Other structures include concretions, mud cracks, bioturbation (destruction of bedding planes by burrowing animals), imbrication (stacking of flat grains at an angle to the bedding plane, with dip upstream) and flute casts (bed gouges dug up by passing turbidity currents), and trace fossils like the worm holes at Dinosaur Ridge (right).
The voids between individual grains in a clastic sedimentary deposit may be reduced during compaction, but they're rarely eliminated. They provide opportunities for cementation of the sediment grains and for fluid transport and storage—issues near and dear to the hearts of petroleum geologists and hydrogeologists.
Porosity is the ratio of residual void volume to total rock volume. In a sandstone, greater angularity and better sorting of the sand grains both increase porosity, while tiny clay grains reduce porosity by collecting in and plugging up the voids. Cements also reduce porosity. Sedimentary rocks with high porosities can host a variety of fluids, from ground water to magmatic fluids to petroleum to natural gas.
Permeability, the ease with which a fluid can move through the pore spaces of a rock, is a different matter, and clay content once again enters the equation. Clay grains have charged surfaces and high surface-to-volume ratios. These properties retard flow by increasing the attractive forces between pore walls and contained fluids.
The stiff, angular quartz grains and the low clay content in a sandstone like the South Platte member of the Dakota Group of eastern Colorado (right) give it high porosity and permeability. Petroleum and natural gas generated in the underlying highly organic Benton shale (the source rock) migrate upward into the South Platte (the reservoir rock), where they are free to flow and collect in folds that serve as traps.
How well consolidated a sedimentary rock becomes depends on
Since cements play a critical role in the color and durability of sedimentary rocks, they deserve a little detail.
Water is by far the most common pore-space fluid found in sedimentary deposits. Water rich in dissolved iron oxides, calcite (calcium carbonate) or silica (silicon dioxide) can deposit these minerals as cement between sediment grains. Well-cemented sedimentary rocks like the Kayenta Formation (right) tend to be resistant to erosion, but if they ever encounter pore fluids capable of dissolving their cements, all bets are off.
Silica (SiO2) is the least soluble and therefore the most durable cement, but it's rare outside sandstones with 90% or higher quartz contents. Hematite (Fe2O3) is another durable cement notable for the wide range of reddish hues it imparts. Calcite (CaCO3) is the most soluble and the least durable of the common cements, particularly in the presence of acidic ground water. When the cement dissolves, the sedimentary rock becomes friable and eventually disaggregates on exposure.
Sedimentary or Metamorphic?
At what point altered sedimentary rocks like those at the Maroon Bells become metamorphic has always been controversial, but there are some widely accepted conventions.
By themselves, end-stage sediments aren't very colorful. Quartz grains are usually clear or gray, most clay minerals are faintly green at best, and carbonates tend to be white, gray or even faintly blue. Yet clastic rocks show an extensive palette of colors, sometimes subtle, sometimes vibrant, including white, gray, black, buff, brown, green, yellow, orange, red and even purple and blue. Where do the colors come from?
Brown, red, orange, yellow and purple hues come from mixtures of iron compounds in pore spaces. Hematite (Fe2O3) is red; iron hydroxide is brown; and iron silicates are green, as are most clay minerals. The iron ultimately comes from weathered ferromagnesian minerals common in igneous and some metamorphic rocks. Hematite cements are common in terrigenous sediments, especially in arid climates, as seen in the Triassic redbeds exposed at right in the north wall of Maroon Creek Canyon in the Elk Range.
The grays through blacks come from free carbon derived from organic materials incorporated into the sediment, as in the coaly layers within the Dakota sandstone at right.
Blue comes from cements containing lawsonite or glaucophane—both high-pressure, low-temperature blueschist minerals found in rocks previously taken to depths of 20 km or more and brought back up again in a subduction zone setting. Some limestones can also take on a bluish cast, but I have no idea why.
Passive continental margins like the Atlantic coast of North America tend to acquire submerged shelves extending tens to hundreds of kilometers offshore to end at the continental slope. The shelves typically consist of flat-lying fine-grained end-stage sediments accumulating to great thicknesses over faulted continental crust. Some shelves, like those bordering the Red Sea, advance out onto oceanic crust as well.
In many ways, continental shelves are the simplest of depositional environments, so we'll start by tracing their constituents through all five phases of sedimentation. These phases occur in all types of sedimentation, but they are most developed in the shelf setting.
Climate strongly controls all five phases, but bear in mind that climate can change drastically between source regions and and depositional environments—as it does, say, between the glaciated alpine peaks and subtropical shores of New Zealand.
Frost-wedging, the most important form of mechanical weathering, depends on the fact that water expands by ~9% on freezing and operates anywhere freeze-thaw cycles repeat through the winter—particularly on rock faces in high mountain regions with wet, temperate climates. When water filling a crack freezes, the expansile force generated is often sufficient to open the crack a bit. If the crack parallels the surface, repeated wedging will eventually split the overlying slab off the rock face to fall into a talus slopes below. Abrasion of bedrock by bedload sediments in a swift mountain stream, or by other rocks embedded in the foot of a glacier, is another common form of mechanical weathering. A final form is pressure release, wherein rocks formed at high pressure split apart spontaneously on dispatch to the surface due to internal stresses. The granite domes of ^Rocky Mountain National Park (right) owe their shapes to the process of exfoliation, which combines pressure release and frost-wedging on a large scale.
Chemical weathering is more effective at dismantling rocks than mechanical weathering by orders of magnitude, but the two are often synergistic. As mechanical weathering reduces clast size, it increases the total surface area susceptible to chemical attack. Conversely, as chemical weathering dissolves its cement, sedimentary rock becomes more easily disaggregated by mechanical weathering processes.
Erosion is the mobilization of weathered rock, usually by a moving fluid, be it creeping glacial ice, flowing water or wind. Erosion feeds weathered rock into the transport system(s) made available by topography and climate. Glacial ice is by far the most effective erosive agent, and wind the least. Short of glaciers, most of the work of erosion is done by bank-full streams and 100-year gully-washers, not by average weather. These floods allow streams to carve imposing canyons without glacial help. A steep gradient and an ample supply of water and bedload helped the Arkansas River cut Royal Gorge (right).
Shelf sediments typically travel great distances to their resting places, weathering all the way. By deposition time, they've usually been reduced to stable end-products—quartz sand, clay mud and dissolved carbonate. But flowing water transports sand, clay and lime quite differently, and the differences lead to sorting.
Since the larger sand particles tend to bump and roll along the bottom until they go back into suspension at silt-size, they tend to settle out first as water velocities diminish in shallow, near-shore settings. The much smaller, lighter, flatter clay particles are carried in suspension and drop out only in much deeper, quieter water. The calcium carbonate is carried in solution and comes out only when actively pulled out by shell-building fresh water or marine organisms like corals and oysters. These light-dependent organisms are usually most abundant in the clear water just beyond the clay deposition zone but under 50 m in depth.
Air masses can spread fine clay dust across oceans, but wind can carry sand only so far up the dune face or down the back side. Dissolved carbonate aside, flowing ice can carry anything it pleases, from mud to house-size boulders. That's why glacial deposits are so poorly sorted.
Note that stream transport can't round quartz grains completely, but surf and wind can come close.
When streams lose their momentum at lower elevations, they drop their loads in roughly this order: Large clasts, small clasts, sand and mud (silt and clay). The lime stays in solution until the water reaches the right ecosystem. This hierarchy effectively sorts sediments, first by size and density, and then, at end-stage, by composition. Thus sandstones are usually near-shore and stream bed deposits, clays are usually flood plain or deeper water deposits, and limestones are deeper still. Mixed sediments—like sandy or limey mudstones (the latter marls), or muddy or limey sandstones, or sandy or muddy limestones—are always possible under the right conditions.
The alterations a sediment undergoes after deposition but short of metamorphosis are lumped under the term diagenesis. Diagenesis is generally taken to operate below 150-200°C, which may come at depths of 3-10 km, depending on the local geothermal gradient. Above those temperatures, metamorphic signatures like foliation and the recrystallization of quartz appear.
Lithification, the formation of rock from sediment, is perhaps the most important outcome of diagenesis. The table below lists the most common diagenetic processes.
Let's follow a shelf sand deposit through diagenesis to lithification. Similar things will happen to other sediments. The sand will compact under their weight of overlying deposits. Compaction increases the surface area of intimate grain-to-grain contact, but not so much with quartz grains, which are rigid and angular and tend to meet only at points. More ductile lithic grains (e.g., of schist or shale) will squeeze into pore spaces to some extent while the stiffer quartz grains hold their ground. The more angular the stiffer particles are, the less the ductile grains can flow. Fluids, most notably ocean water, ground water or petroleum, will circulate through the residual pore spaces, depositing any over-saturated minerals they might contain as cement. The cement increases the strength of the forming rock by further increasing the total surface area of effective grain-to-grain contact. At any point in the process, the residual pore spaces can hold fluids which may add or dissolve cement.
Heating (up to 150°C) is usually the most important agent in diagenesis. Increased pressure due to burial usually acts only through compaction. Pressure-related changes in mineralogy occur only in metamorphic settings.
Shelf sediments seldom reach the deep ocean floor of the abyssal plains. Transport mechanisms are generally lacking, and sinking carbonate skeletons dissolve before reaching the bottom.
In pelagic (deep open-ocean) settings below the carbonate compensation depth, radiolarian cherts dominate the sedimentary veneer covering the igneous rocks of the oceanic crust. Insoluble silica-based radiolarian skeletons rain down from the sunlit surface layer onto the ocean floor to make ~1 mm of chert per 10 Ka at any depth. With burial and heating, the skeletons recrystallize to form thin (10-50 mm) beds of nearly pure microcrystalline quartz separated by even thinner beds of shale chemically segregated from the silica during recrystallization. The shaley partings derive from microscopic clay dust particles dropped over the oceans by global winds passing over the continents. Such sequences are also known as pelagic, bedded, or ribbon cherts. No other sediment sources are available to the oceanic crust of the abyssal plains.
In the field, pelagic cherts appear as bedded or ribbon chert, sometimes in layers well over 50 m thick. The sporadic horizons of small chert nodules found in limestones like the Paleozoic limestones of western Utah around Canyonlands National Park represent chemical segregations of minor inputs of radiolarian silica in a shallow biogenic depositional environment dominated by carbonate inputs. Any chert found in limestone must have been deposited above the carbonate compensation depth.
As a transgressing sea like the Cretaceous Interior Seaway invades a landscape, it first lays down beach and barrier bar deposits. As it deepens, it deposits clay over the sands. As it deepens further, carbonates accumulate on the bottom as corals and shellfish set up shop in the clear water. If the sea regresses, muds bury the limes, then sands bury the muds as the water shallows. Terrigenous (dry land) strata then bury the entire marine sequence once the water is gone.
On burial and lithification, such depth-sorted sediments form a sandstone-shale-limestone-shale-sandstone sequence marking the coming and going of the sea. A Cretaceous to Paleocene transgressive-regressive sedimentary sequence crops out throughout Colorado, notably at the Dakota Hogback (above right) west of Denver. There, the marine strata are, from oldest to youngest, are the transgressive Dakota sandstone, the Benton shale, the Niobara limestone, the Pierre shale, and the regressive Fox Hills and Laramie sandstones. The terrigenous Arapahoe and Denver Formations still cover the younger portions of the sequence. The Niobara hogback is long lost to quarrying.
The marine deposits we've examined so far make up a substantial fraction of the sedimentary rocks found in Colorado, but terrigenous sediments deposited on dry land are also strongly represented. With few exceptions, the voluminous Late Paleozoic Fountain, Maroon and Hermosa sediments shed from the Ancestral Rocky Mountains, most of the pre-Dakota Mesozoic strata, and the sediments shed from the current Rockies during and after the Laramide orogeny were transported and deposited on land by streams. Lacustrine (lake bed) deposits like the Eocene Green River Formation are transitional in this regard.
The table below lists some important terrigenous depositional environments, derivative Colorado rocks and modern examples. Formations listed together accumulated in laterally equivalent or persisting environments; those listed under separate bullets represent recurrences of the same depositional environment.
Hematite cements released by the weathering of ferromagnesian source rocks grace many terrigenous deposits with a range of reddish hues—hence the term redbeds. An arid depositional climate seems to help bring out the red, and such environments were the rule rather than the exception in Late Pennsylvanian through mid-Jurassic Colorado. At right, Late Paleozoic redbeds of the Fountain Formation crop out in Red Rocks Park near Denver. The photo at the top of this subsection shows Triassic redbeds exposed in Maroon Creek Canyon in the Elk Range. Marine sediments are less likely to show such colors.
In keeping with the worldwide dominance of clastic rocks at the surface, most of the bedrock units shown on the Geologic Highway Map of Colorado are largely or entirely composed of clastic sedimentary rocks, particularly the Cenozoic strata. Colorado is under quota in this regard, however, because erosion has already stripped the sedimentary cover from the summits of many of its Laramide uplifts.
Skip to Metamorphic Rocks
Sedimentary rocks precipitating from an over-concentrated solution (say, from the dregs of an evaporating sea) or manufactured by organisms are said to be chemical, and the latter, biochemical. They're generally quite different than their clastic brethren and constitute less than 5% of all sedimentary rocks. Almost always water-laid, chemical rocks supply their own cement to any water able to circulate through their pore spacess, but they're generally low in both porosity and permeability. Petroleum geologists don't go looking for oil in chemical rock formations, but chemical rocks can seal off petroleum traps in underlying strata.
Chemical rocks typically weather by dissolving into surface or ground waters with or without the aid of mechanical weathering. Ground water is more often than not acidified by organic materials (less often by acid rain), and that makes it a highly effective solvent for chemical rocks, especially the carbonates. Given their solubility, chemical sedimentary rocks are understandably much less common in outcrop than their clastic counterparts, but chert, a tough chemical rock made of very stable microcrystalline quartz, is one of the most resistant of all surface rocks, and limestones can form spectacular landscapes in the right setting—e.g., the famous ^karst towers of Guilin, Guanxi Province, south China.
Carbonate rocks are by far the most important of the chemical rocks—both by volume and by virtue of their critical role in the planet's carbonate cycle, which among other things, controls the carbon dioxide content and therefore the green-house properties of the atmosphere. Most carbonate rocks are biochemical.
Limestone forms directly from the compacted calcium carbonate skeletons and shells of corals and other calcareous organisms. It takes many, many different forms, depending primarily on the depositional environment. Physicochemical processes related primarily to decreasing temperature and increasing dissolved CO2 concentration with depth in ocean water force precipitated carbonate minerals back into solution below the carbonate compensation depth (CCD), which today is ~4 km down. Since marine limestones can only form above the CCD, geologists find them a useful if rather coarse paleodepth indicator.
Dolomite, another important biochemical carbonate rock, is a chemical variant of limestone in which some of the calcium atoms are replace with magnesium. Interestingly, in Paleozoic and earlier times, dolomites outmassed limestones by about four to one, but dolomite comprises only a tiny fraction of the carbonate rocks forming today. No one knows why. Despite the name, the Mississippian Leadville Limestone of Colorado mining fame is mostly dolomite.
Marls are mixtures of clays and carbonates with at most minor amounts of quartz. Weathered marls make challenging soils for growers and civil engineers alike.
Cherts are very hard, highly resistant and often colorful chemical rocks of microcrystalline silica that tend to form topographic highs wherever they crop out in bedded form. Weathered bedded chert can look like a fine-grained sandstone from a distance, but fresh surfaces are lustrous and slightly translucent. Chert colors range from pure white, gray and black to yellow, green, rust and red, depending on the impurities involved. Chert nodules ranging up to meters across and weathered or pried from Paleozoic limestones have been an important raw material for the native stone toolmakers of the US Southwest, who valued chert's exceptional hardness and conchoidal fracture habit. Pelagic cherts provide geologists with reliable markers of deep ocean floor environments.
Most chert forms indirectly from planktonic debris—specifically, from the accumulated siliceous (silica-based) skeletons of radiolaria and diatoms after recrystallization of their silica on burial heating to temperatures of 80-100°C. [??] Logically, chert formed in this manner might be classified as biochemical, but because of the required recrystallization step, this transitional rock type is classed as chemical by convention. Pelagic radiolarian cherts form on the ocean flow below the carbonate compensation depth. Diatomaceous ribbon cherts form in carbonate-starved lake settings under certain conditions.
In shelf environments, cherts usually appear as small nodules chemically segregated and recrystallized from the minor biogenic silica trapped in massive limestones. For reasons that remain unclear, the nodules tend to line up along limestone bedding planes.
Truly chemical or inorganic cherts do form occasionally in lakes that become severely alkalinized on a seasonal basis—e.g., by blooms of photosynthetic algae or by winter drainage from nearby sodium carbonate lavas (carbonatite). Normally undersaturated dissolved silica becomes oversaturated at high pH (> 10) and precipitates out as a silica gel that crystallizes directly to chert with burial. Colorado's Morrison and Green River Formations contain inorganic cherts.
Evaporites are perhaps the purest form of chemical rock. They typically appear in the rock column as massive gypsum (CaSO4) and rock salt (halite, NaCl) deposits left behind by an over-saturated body of natural water trapped in a restricted basin with a limited refresh rate. Evaporites usually form in arid climates with a high evaporation rate and little rainfall. ( Think Bonneville Salt Flats.) In the geologic record, they've most often formed in the planet's desert belts at 30-40° paleolatitudes.
Since evaporites are highly soluble and easily eroded, they're not often found at the surface, but the razor-sharp halite foothills of the west slope of the Zagros Mountains of Iran are a notable exception. Gypsum is the least soluble of the common evaporite minerals and therefore the most likely to be found in outcrop.
Late Paleozoic evaporites of Ancestral Rocky Mountain provenance are common in the Maroon and Paradox Basins of central and southwestern Colorado, respectively. Due to their unusual chemical and mechanical properties, evaporites make for odd topography, like the chaotic Pennsylvanian black shales disrupted by rising gypsum bodies along the Colorado River between Dowd and Wolcott on I-70.
From both an outcrop and economic standpoint, one of the most important chemical sedimentary units in Colorado is the Mississippian Leadville Limestone lumped into the dark purple Early Paleozoic "MDOC" unit on the Geologic Highway Map of Colorado. The thick evaporites of the Maroon and Paradox Basins are mapped under the light blue "PP" unit. The inorganic cherts of the Green River and Wasatch Formations will be found within the pale yellow "Tl" units of western Colorado.
Skip to Rock Formations
Igneous and sedimentary rocks are stable only at the conditions under which they became rocks. Since most rocks form at some depth below the surface, their equilibrium temperatures (T) and pressures (P) are usually higher than those found at the surface—sometimes much higher. If uplift and erosion return a high-T, high-P granite to the low-T, low-P surface, it will adjust its new environment via chemical and physical weathering, as we saw in the discussion of sedimentary rocks above.
But what about the opposite scenario? If a continent-continent collision deeply buries a shale stable under low-T, low-P near-surface conditions, its elements will reshuffle into a new mineral assemblage better suited to the high-T, high-P conditions found at depth. Along with the chemical adjustments come textural changes stemming from recrystallization along preferred orientations, microfracturing, and even plastic flow. These alterations go beyond diagenesis to enter the realm of metamorphism. The shale, the parent rock, can give rise to many different metamorphic products, depending on the conditions involved. If those products return to the surface, they can yield important clues regarding where the rock's been and what kind of rock it might have been when it started its journey.
Lest you think that metamorphism is just something that happens to rocks that stray too deep, consider the story of a certain high-pressure garnet-bearing metamorphic rock called eclogite. The Ligurian Alps and Ring Mountain in northern California are among a handful of places in the world offering substantial in situ exposures. Eclogite isn't rare—by now, the upper mantle is riddled with the stuff. It just takes an extraordinary upheaval to bring it to the surface.
But all mantle solids are rare at the surface. What's so special about eclogite? Simply this: Subduction drives all of plate tectonics, and eclogite is its whip. In fact, a critical repositioning of the depth of eclogite formation by the cooling of the earth was an important precursor to the onset of plate tectonics at ~2.0 Ga.
Continuing top-down cooling of the upper mantle produces gravitationally unstable oceanic lithosphere that progressively cools, thickens, stiffens, imbibes water and becomes more dense through cooling and phase changes as it rafts away from its spreading center of origin. Not far from its spreading center, the oceanic lithosphere (slab for short) is already denser and stiffer than the hot asthenosphere on which it rests. Flexural support and lateral compression by surrounding lithosphere allow it remain topside for a while despite the density inversion, but eventually, all overdense slabs buckle downward and sink subvertically into the upper mantle along convergent plate boundaries known as subduction zones. The rate at which gravity pulls down the slabs sets the pace for the rest of the system—continental drift, arc magmatism, the works.
That's where eclogite comes in. Once the falling slab reaches a depth of ~60 km, the basalt and gabbro making up its crust rapidly convert to much denser eclogite—denser even than the upper mantle rock below the asthenosphere. Like a lead sinker on a fishing line, the eclogite hastens the slab's fall. Moreover, water released into mantle wedge overlying the slab as a chemical byproduct of the eclogite transition eventually triggers the partial melting that fuels arc magmatism, an important pathway for the cooling of the earth and the recycling of upper mantle into continental lithosphere.
So eclogite helps the world churn and burn. It'd be a much duller place without the metamorphism that produces eclogite.
To make the metamorphic process less abstract, and less exotic, let's follow a common shale through its metamorphosis as both temperature and pressure progressively increase. In this example, shale is the unaltered starting point or parent rock for the metamorphic process.
Table Note: The external links lead to illustrated descriptions of metamorphic rocks at Lynn Fichter's Metamorphic Rocks site.
Under increasing temperature and pressure, a mature quartz sandstone would have evolved quite differently than the shale followed through the regional metamorphic process in the table above. After recrystallizing into a ^quartzite at relatively low temperature and pressure, it would have resisted all further changes short of melting. In the field, quartzites and well-cemented pure quartz sands can be hard to tell apart, but a hammer and a hand lens can help: Fractured quartzites generally break through grains, sandstones between grains.
When alterations due to increased pressure-temperature (PT) conditions or permeating fluids exceed those associated with diagenesis, the rock becomes metamorphic. The higher the temperature or pressure, the higher the grade of metamorphism.
Practically speaking, the boundary between diagenesis and metamorphism is hard to define. Since elevated temperatures are much more effective than elevated pressures in altering rocks early on, it's common to define the lower limit of metamorphism in terms of a temperature threshold, usually 150°C. Under typical geothermal gradients, the temperature reaches 150°C at a depth of 3-4 km and a pressure of 0.5-1.0 k bar.
The upper limit of metamorphism is the onset of significant melting, which occurs at 650-800°C for most crustal rocks. Such temperatures are common in the lower crust near the Moho (the crust-mantle boundary) but can also occur in the upper crust in conjunction with magmatism. The presence of water, whether free in pore spaces or bound in hydrated minerals, significantly lowers the PT conditions needed to melt nearly any rock. Water is scarce at lower crustal levels, where pore space is scant and anhydrous minerals are favored, but it becomes progressively more abundant at middle and upper crustal levels.
Partial melting under extreme PT conditions is called anatexis. It's more often the result of tectonism (regional metamorphism) than magmatism (contact metamorphism) because magmas don't generally carry enough excess heat to melt significant quantities of surrounding rock and remain molten at the same time.
Migmatites, rocks with both igneous and metamorphic features, are fairly good upper limit markers in the field. They always reflect extreme PT conditions, but not all migmatites are anatectic.
Temperature and pressure usually rise together during regional mountain-building, the most common cause of metamorphism. The same is true of deep burial in a subsiding sedimentary basin like the Denver basin on the east flank of the Rockies.
But other telling PT combinations are observed. Normal subduction zones, for example, offer unique high-P, low-T environments where otherwise rare blueschists containing minerals like lawsonite and glaucophane can form. In the field, blueschists serve as reliable signatures for subduction tectonics.
Igneous intrusions, on the other hand, tend to create high-T, low-P metamorphic conditions in surrounding country rocks, which can also be altered by permeating magmatic fluids. In the field, the characteristic alterations can point to an otherwise unsuspected intrusion nearby. Closing backarc basins at collisional subduction zones represent a special case here. These very hot orogens contain high-T, low-P environments capable of generating granulites.
Rock go metamorphic for many reasons. Mountain-building is the most common cause, but igneous intrusions, faulting, deep burial in sedimentary basins and a variety of ocean floor processes are also effective in altering rocks beyond the limits of diagenesis. At rock level, pressure, heat, and fluids do the actual work of metamorphism.
Pressure increases come with sedimentary and tectonic burial and with other tectonic stresses. Elevated pressures foster chemical change by forcing pre-existing grains into more intimate contact and by favoring the formation of denser reaction products. Pressure gradients prod crystals with elongate (rod- or disc-like) forms to regrow preferentially in directions perpendicular to the gradient. High partial pressures of volatiles like water and carbon dioxide can influence the course of metamorphic reactions and, in the case of water, can reduce the melting points of most rocks. Elevated lithostatic pressures also reduce melting points, with or without water.
Temperature increases with burial, sedimentary or tectonic, and also with nearby magmatism. Rising temperatures encourage chemical adjustments by destabilizing existing minerals and by accelerating the diffusion of reactants. Elevated temperatures also favor the formation of refractory minerals. At high temperatures, grain boundaries dissolve, cracks "heal", previously segregated minerals come together and react, both in solid phase and through fluid phase intermediaries, and new products crystallize out in residual pore spaces.
The overall effect metamorphic heating is to form new grains of larger and roughly equal (equigranular) size with or without chemical change. In the presence of a pressure gradient, for example, high temperatures mobilize mineral molecules at high-stress contact points and allow them to recrystallize in low-stress pore spaces, effectively reorienting the crystals as the gradient dictates without altering their chemistry. Temperature-driven isochemical recrystallizations tend to dominate metamorphism at low grades, as in the transformation of shale to slate.
Fluids enter the fray via groundwater, seawater, magmatic emanations and the release of volatiles. Depending on the PT field involved, pore-space fluids can exist in liquid or vapor phases and can influence chemical and physical reactions in a number of ways. Water transforms anhydrous into hydrous mineral forms and lowers the melting point of the former even if hydration reactions don't occur. High pressures make volatiles more reactive by increasing their partial pressures. High temperatures make everything more reactive.
Metamorphic products vary with the PT field applied and with the presence of fluids. Below are the most common types.
Earth processes leading to metamorphism tend to involve large areas. Compressive mountain-building is the prime example, but subduction zones and regional magmatism bring metamorphic change to large regions as well. Regional metamorphism typically brings elevations in both temperature and pressure, sometimes to extremes. Aside from the loss of volatiles, regional metamorphism is usually isochemical, meaning that the metamorphic rocks and their parents retain very similar elemental compositions. However, the products of regional metamorphism tend to become more and more granite-like at high grades. This type of metamorphism is also known as Barrovian after the pioneering British metamorphic geologist George Barrow.
Contact metamorphism is a high-T, low-P alteration of the country rocks surrounding an igneous intrusion. A depth of 6 km and a pressure of < 1.5 kbar would be a common scenario. Hardened, reddened (oxidized) country rocks mark the thermal aureole or contact zone ("bake zone"), which can extend for meters to kilometers around the intrusion.
In silicate country rocks, the resulting metamorphic rock a hornfels. With carbonate country rocks, a marble or skarn. As with most forms of high-T metamorphism, grain sizes usually increase. At the low confining pressures typically involved, hornfels grains develop little if any parallelism—indeed, any pre-existing parallelism may be lost in contact metamorphism. The photo at right shows a distinct gradient in the contact metamorphism of red Maroon Formation rocks to green hornfels near Cathedral Lake, Elk Range.
If hot, mineral-laden fluids emanating from or mobilized by the intrusion further alter the country rocks, the process is called metasomatism. Metasomatism is single-handedly responsible for the mineralization of the Colorado Mineral Belt. Metasomatic changes are usually most pronounced when carbonate rocks surround silica-rich intrusions, and that's precisely the combination—silicic Laramide and mid-Tertiary intrusions placed within or near the Leadville Limestone—that set Leadville and Aspen on the road to mining fame. The metasomatic aureole surrounding an intrusion is usually less extensive than the thermal aureole.
Another form of high-T, low-P granulite-fascies metamorphism occurs during closure of backarc basins, which are heated both by arc magmatism and by hot asthenosphere brought near by lithospheric thinning beneath the basins. This backarc basin signature is valuable to geologists investigating collisional continental crust.
Sedimentary rocks like graywackes caught up in the accretionary wedge of a subduction zone and entrained by the cold sinking slab can quickly reach depths of 20 km or more. If they're spit back out before they can heat up to temperatures appropriate for such depths, they reappear as blueschists. (No one knows exactly how blueschists rise again, but buoyancy relative to surrounding mantle is probably involved.) The amphibole mineral glaucophane provides the blue cast.
No other known tectonic process can produce the high-P, low-T "blueschist" minerals (lawsonite, glaucophane, jadeite, pumpellyite) that develop in the subduction environment. Finding blueschist in the field is a surefire marker of subduction.
Eclogite is another high-P, low-T metamorphic product of subduction formed at ~60 km by a phase change in the basalt and gabbro making up oceanic crust. Eclogite's great density and depth of formation make it an important subduction accelerant. It returns to the surface only under very unusual circumstances.
Hydrothermal brines circulating through abundant cracks in the newly-formed oceanic crust near spreading centers chemically alter the seafloor basalts and gradually fill the cracks with calcite precipitated directly from the hot seawater. The resulting green fine-grained igneous rocks, rife with crisscrossing white veins, are called greenstones. They find their way onto land in subduction zone collisions that uplift fragments of oceanic lithosphere or accretionary wedges. The Early Proterozoic Green Mountain arc and Farwell-Lester shear zone of northern Colorado record such collisions during the early assembly of the Colorado Province.
Cataclasis results when rocks around faults or other intense deformations fracture on very small scales. Reduction in grain size is the most common signature; metamorphic minerals don't usually form because temperatures are typically low. At greater depths, mylonites form by recrystallization with crystal orientations dictated by local stresses or plastic flow. At right is a 1.7 Ga quartz monzonite mylonite exposed in Golden Gate Canyon State Park in the Front Range.
Recrystallized mylonites can end up stronger than the undeformed rocks around them. Vertical fins of tough 1.7 Ga mylonite (right) mark the Homestake shear zone (SZ) along Homestake Creek in the northern Sawatch Range. The cataclastic foliation is obvious even at this distance.
When sediments interlayer with volcanic rocks in a deep sedimentary basin like the Denver basin, they can reach temperatures of several hundred degrees. High water pressures foster changes in mineralogy without disturbing rock textures or fabrics. Zeolites are a common product of burial metamorphism, but some geologists consider it to be nothing more than high-grade diagenesis.
Since chemical and physical metamorphic reactions rarely take more than days to weeks to complete in the laboratory, even in solid rock, they have no trouble keeping up with changing pressure-temperature (PT) conditions in metamorphic environments, where PT conditions typically vary on timescales measured in thousand to millions of years.
With rocks that responsive, you might think that they'd also readily revert to their original state if the PT changes were undone, but that's seldom the case. Chemical and textural metamorphic alterations are generally irreversible. Once a shale progrades into a gneiss (like the Royal Mountain hornblende gneiss at right), stable at, say, 500°C, it can't retrograde back to a shale on uplift and re-exposure at the surface. It will weather and erode as a gneiss instead. Granted, the feldspars and ferromagnesian minerals in the gneiss may well break down into clay redeposited as shale one day, but a direct reversion to shale can't happen.
Among the many reasons for irreversibility are the general increase in grain size, the decrease in porosity and the release of volatiles (most notably water and carbon dioxide) common to many metamorphic alterations. Once the volatiles escape the rock, there's no turning back—at least not to the pre-metamorphic mineral assemblage. Retrograde metamorphic minerals may form in very small quantities (for example, at the margins of higher grade crystals), but prograde minerals dominate any metamorphic rock returned to the surface. If the minerals don't revert, the metamorphic fabrics and texture they create won't change, either.
Note that the irreversibility of metamorphic changes means that a metamorphic rock faithfully records in its mineral assemblage and fabrics the maximum PT field it's endured. That in turn tells a good bit about where the rock's been. Geochemists and geophysicists have been busy working out which metamorphic changes go with which PT fields since the Finnish geologist Eskola first proposed such associations in 1915. Geologists seeking to unravel metamorphic terrains make heavy use of such relationships.
In the field, metamorphic rocks are often identified as such by their distinctive fabrics and textures. What's the difference? Fabrics relate to the alignment of grains, while textures relate to the distribution of grain sizes and and the normalcy of grain shapes. Metamorphic fabrics and textures usually replace similar pre-metamorphic features, like flow-related sedimentary grain alignments, but they often superimpose in previously metamorphosed rocks.
Metamorphic minerals with elongated or platy shapes respond to unequal stresses by recrystallizing in orientations that reduce strain. That usually means putting the new shortest grain dimension along the axis of greatest compression. A grain alignment produced in this way is called a metamorphic fabric—one of hallmarks of metamorphic rock. If you find a crystalline rock with a fabric, it's most likely metamorphic, but note that non-metamorphic intrusives can acquire flow fabrics through magma currents present prior to freezing.
Fabrics can be dated by noting which independently datable structures (like plutons) they do and do not involve. Features like wings on porphyroclasts can be used to determine their sense of motion. Rocks with complex deformation histories often display several superimposed metamorphic fabrics; teasing these apart in the field is one of the goals of structural geology.
Foliation arises from the coplanar alignment of platy crystals. Strong foliation involving small grains results in cleavage—the ability to split a rock in to highly planar sheets. Schistosity is the alignment of coarse platy or elongated grains. Schistose rocks can still be broken into small planar flakes, but not into large sheets. Subsequent stresses can deform the plane of schistosity, as seen in the porphyroblastic schist at right from Golden Gate Canyon State Park.
Lineation arises from the alignment of elongate but not platy grains, or from the intersection of two non-coplanar foliations. Note that lineations are not just foliations seen on end. They are truly one-dimensional fabrics, but most lineations turn out to reflect alignments of elongated grains within a plane of foliation.
Gneissic banding reflects the segregation of felsic (light-colored) and mafic (dark-colored) minerals perpendicular to the maximum compressive stress. The bands may range from millimeters to several meters in thickness and may be straight or highly contorted, particularly is the rock's undergone plastic flow.
Whereas fabrics relate to grain alignments, textures relate to the distribution of grain sizes and the normalcy of grain shapes. Metamorphic processes generally foster larger grains of roughly equal size. A metamorphic rock with such an equigranular texture is said to be granoblastic. That said, porphyroblasts, crystals much larger than average grain size, are not uncommon in metamorphic rocks, which are then said to be porphyroblastic in texture.
Cataclastic (high-strain) metamorphism produces mylonites. Under high-strain conditions, grains of particular minerals (often feldspars) can aggregate into prominent porphyroclasts embedded in a much finer foliated matrix. If the porphyroclasts are eye-shaped, as they often are, they're called augen. In the field, certain porphyroclast shapes serve as valuable sense-of-shear indicators in porphyroclastic metamorphic rocks. Veins may be pinched and pulled into sausage-shaped bodies called boudins.; the process is called boudinage. In the tortured schist at right, the light-colored pegmatite veins show boudinage.
Under ideal conditions, minerals grow into crystals with characteristic shapes, but when new minerals grow or old ones recrystallize in a real metamorphic rock, they may or may not be able to assume their ideal shapes. If they succeed, the rock is said to have an idioblastic texture. If they fail badly, the texture is xenoblastic (strange). In between are hypidioblastic textures.
Wavy strands and sausage-shaped boudins generated by plastic flow under extreme PT conditions are seen in migmatites at the brink of wholesale melting and in ductile banded gneisses. Such features are transitional between fabrics and textures. The wavy white feldspar strands within the Homestake shear zone mylonite at right indicate deformation at temperatures near the melting point; sense of shear can be inferred from the wavy shapes.
For more on metamorphic textures and fabrics see Lynn Fichter's ^Foliated Textures of Metamorphic Rock.
The minerals produced by metamorphism are determined to some extent by parent rock composition but to a greater extent by the maximum temperature and pressure to which the minerals equilibrated, with little if any retrograde alteration on return to the surface. Certain metamorphic mineral assemblages have thus become associated with certain PT fields. These associations are called facies. The table below gives an idea of the minerals and PT fields involved; the temperatures and pressures shown are very rough approximations.
Identification of the mineral facies present in a metamorphic outcrop thus gives an idea of the maximum PT conditions, and to a lesser extent, of the parent rocks involved. When a rock contains blueschist facies minerals, for example, it's said to be of blueschist grade, and we cans safely infer it's spent time in a high-P, low-T subduction zone. Other facies are less specific with regard to the tectonic environment but provide important information nonetheless.
The metamorphics derived from quartz sandstones (quartzites) and mudstones (slate, phyllite, schist, gneiss and migmatite) have already been discussed in some detail above. Note, however, that schist, gneiss and migmatite are final common metamorphic pathways for many parent rocks, not just mudrocks. This section describes some other common metamorphic rocks. 
Table Note: The external links lead to illustrated descriptions of metamorphic rocks at Lynn Fichter's Metamorphic Rocks site.
On the Geologic Highway Map of Colorado, the 1.7 Ga metasediments and metavolcanics making up much of the Precambrian basement are mapped as "Xm" in medium gray, while the 0.9-1.4 Ga metasediments of the Uinta Group and the Uncompahghre Formation appear in medium gray with wavy lines under the label "Ym". The craggy hornfels outcrops of the southern Elk Range are lumped with the Maroon Formation under the light blue "PP" unit.
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In this section, we'll discuss some of the basics of stratigraphy, the science of deciphering stacks of sedimentary rocks.
A formation is a set of rock layers that share an age range and one or more distinctive physical characteristics, can be followed for some distance as a group, and can be drawn on an outcrop map. The more recognizable layers within a formation are often called members or units. Since formations deal in layered rocks, we usually think of sedimentary formations, but it makes just as much sense to speak of volcanic formations in some settings, and minor ash units are tolerated in predominantly sedimentary formations like the Cretaceous Pierre Shale. At right, the Grand Monocline at Colorado National Monument exposes the distinctive and widespread Morrison Formation.
The mappability criterion is an important one, for the true extent and distribution of a sedimentary package yields important clues about its provenance. Many of the synorogenic sedimentary strata of the upper Denver Basin vary too much in thickness, lateral extent or even sediment type (fascies) to be truly mappable. When Robert Raynolds undertook to reorganize the confusing welter of named strata (some already with formation status) into mappable units, he ended up with only two — an older D1 sequence and a younger D2 sequence separated by an easily recognizable unconformity and paleosol. Unconformities also demarcate the bottom of the D1 and the top of the D2.
Formations consisting mostly of a single rock type receive a place name indicating the type locality (the place where the formation was first recognized) followed by the name of the dominant rock, as in "Leadville Limestone". Formations featuring more than one significant rock type get a place name followed by "formation"—hence the "Morrison Formation", which contains claystone, sandstone and minor limestone members. Formations found together over large areas may be lumped into a group. The Dakota Group of eastern Colorado subsumes an upper South Platte Sandstone and a lower Lytle Formation member.
Formations are often divided by terms like upper and lower to reflect the relative positions of included rock units as they appear in the field. The upper units are usually younger than the lower units, but a word of caution is in order here: Things can get upside down now and then. In a sedimentary formation caught up in an overturned anticline, the lowest units will be younger than the middle strata.
Upper and lower formation boundaries may be fairly distinct on paper, but they aren't always conspicuous in the field. For instance, the lower sands of the early Cretaceous Lytle Formation are distinguished from the upper sands of the late Jurassic Morrison Formation only by the presence of small chert pebbles weathered out of Paleozoic limestones and dolomites exposed far to the west. You'll have to get pretty close to tell them apart.
Unconformities (depositional gaps wherein one formation accumulates on the eroded surface of another) usually provide the sharpest and most conspicuous formation boundaries, especially if the strata change dip across the gap — in which case it's an angular unconformity. Proponents of sequence stratigraphy would argue that unconformities are also the most useful of stratigraphic boundaries.
At right, gently dipping late Devonian sandstone of the Elbert Formation rests unconformably on nearly vertical 1.6 Ga Uncompahgre Formation quartzites above Box Canyon Falls just south of Ouray off US 550. This famous angular unconformity draws geologists from all over the world, but the volunteer working the visitors' center knew absolutely nothing about it. Why am I not surprised?
Formations have a nasty habit of changing laterally in devious ways. Eventually, they all die out against the margins of their receiving basins, some sooner than others. In one form or another, the upper Cretaceous Dakota Group has a 9-state distribution, but the synorogenic Laramide sediments east of the Front Range cover at most an 80 x 150 km area along the mountain front. They get thicker here and thinner there and pinch out altogether at many points in between.
Formations may have a distinct lateral margin in one location (e.g., against an erosional surface or a fault, as in the photo at right) and interfinger with laterally equivalent formations in messy ways elsewhere, or they may just up and change names at some arbitrary and usually unspecified boundary for reasons that are almost always more historical (in the history of science sense) than logical. Thus we're left with a Dakota Sandstone in western Colorado depositionally equivalent to the upper South Platte Sandstone of the Dakota Group of eastern Colorado. Formations may also contain lenses of minor or even major units here but not there. What a headache!
Such is the inherently untidy nature of sedimentary and volcanic transport and deposition in the real world. Throw in some faulting and a small uplift or intrusion here and there, and you can easily end up with a huge mess in the field.
When dealing with rocks and stories about rocks, one must bear in mind at all times that "formations" are really nothing more than hopeful human constructs foisted upon a planet with no apparent limit to its taste for complexity. Nevertheless, formations give us a convenient and often useful if flawed way to speak of and think about earth materials with common histories. They help us chunk certain details long enough to work out a bigger picture of things in time and space. That's how they earn their keep.
Recent advances in stratigraphy now favor sequences over formations as the logical and map units of choice for many purposes, including the reconstruction of paleogeographic settings and the events that punctuate them. A sequence is a stack of related sediments bounded above and below by an unconformity or an equivalent updip conformity. Event-oriented sequences mark distinct depositional regimes punctuated by tectonic rearrangements of depositional basins, source areas or both. Sequence stratigraphy's seemingly endless terminology can be off-putting at first, but it's carefully crafted to bring the full power of ^Steno's principles (first formulated in the late 1600s) and ^Walther's law to bear in the pursuit of depositional origins. The table below summarizes these tried and true maxims of sedimentology.
Raynolds' D1 sequence of synorogenic strata east of the Front Range tells of the erosional unroofing of a rising central Front Range basement block covered at the time (~67-68 Ma, in the Paleocene) with Laramide-aged andesitic volcanics. As this depositional event came to a close, erosion attacking the top of the D1 created an unconformity. Some 8 Ma later, the arrival of the D2 sequence heralded the rise and accelerated erosion of the Pikes Peak and Rampart Range segments of the Front Range to the south. These new source areas shed feldspar-rich (arkose) gravels, sands and muds derived mostly from the Pikes Peak granite into the same basin that trapped the D1 sequence.
Here's a decent ^primer on sequence stratigraphy. While most sequence stratigraphy sites focus on marine sediments, it's no less applicable and profitable in terrigenous environments.
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The idea of bending and flowing solid rock flies in the face of common experience, but once you take it to heart, the Colorado landscape becomes much easier to understand. Here, we'll try to make the concept a bit more believable.
Take a cold stick of butter from the refrigerator. It's solid yet plastic. Press on it with the back of a teaspoon, and it flows slowly out the way but retains its deformed shape when you remove the spoon. Whack the other end of the stick with the spoon. Note that the butter flows more when you apply the deforming force over time. Now scrape the stick with the edge of your spoon and note how it erodes.
Next, heat the stick in the microwave until it visibly softens short of melting. Let it cool a minute or so, then press and scrape again with the spoon. The butter flows and erodes much more readily but still retains its shape when you remove the spoon. (If it doesn't, you partially melted it.) After a night in the freezer, the butter would hardly yield to your spoon press or scrape, but you might be able to break off a piece with a really good whack.
Believe it or not, rock behaves a lot like butter. Sure, the forces, temperatures and timescales involved are orders of magnitude greater with rock, but the basic features of its plastic behavior are much the same. Its plasticity increases dramatically with temperature, especially near its melting point. At low temperatures, like those at the earth's surface, it becomes brittle enough that it would more likely fracture than flow if deformed rapidly.
Despite clear geologic evidence to the contrary, our usually reliable day-to-day sensibilities maintain that rock neither bends nor wears. The disconnect occurs because, unlike butter, solid rock reveals its malleability only over time scales very long compared to human events—typically over spans of tens of thousands if not millions of years. We can barely conceive of such intervals, but given that kind of time, rock will yield to flowing water and will bend or flow rather than break.
Even if it's beyond our ken, the planet's had ample time for the bending and wearing of rock—hundreds to thousands of times over, in fact. Cold sedimentary strata near the surface can be folded without breaking in under a million years. It takes a properly motivated mountain river only a few times longer to cut a 2,700' deep hard rock canyon like the Black Canyon of the Gunnison. It took only 50 Ma to uplift and erode away the Ancestral Rocky Mountains and to bury half the state in their remains. You could have done that 40 times just since the onset of plate tectonics at ~2.0 Ga.
Under great confining pressures, or at depths where temperatures along the geothermal gradient where they reach a significant fraction of their melting points (typically 10-15 km), rocks that are quite brittle at the surface become sufficiently plastic to deform without fracture at rates comparable to the rate at which fingernails grow ( ~10 mm/yr, or 10 km/Ma). Granted, that makes molasses look downright mercurial, but then relative viscosity is the whole idea here. My piano tuner's job security rests on the fact that properly tuned piano wire flows (stretches) at a similar rate and falls out of tune in a matter of months. Over a few Ma, fingernail speed is plenty fast enough to fold great thicknesses of sedimentary rock over the east edge of the relatively brittle Front Range basement block, as in the photo at right. Over 50 Ma, an entire mountain range like the Ancestral Rockies can rise up and vanish.
At right is a 1.7 Ga mylonite exposed at Hornsilver Ridge along the Homestake shear zone in the northern Sawatch Range. This specimen deformed some 10-14 km below the surface. The wavy pink feldspar bands are clear signs of highly plastic flow. Mylonites develop their metamorphic textures at depth near moving faults, which add frictional heating to the mix. This one came close to melting but still deformed as a solid.
The dark Z-shaped inclusions in the final image at right are not the work of Zorro. They are recumbent isoclinal folds in 1.7 Ga schist from the Idaho Springs-Ralston shear zone at Golden Gate Canyon State Park. Distinct folds like these can only form in solid rock with considerable plasticity. Molten rock would never preserve such structures.
Local geothermal and pressure gradients interact with the materials in the local rock column to establish one or more brittle-ductile transitions (BDTs) within each distinct column in the crust. Rocks above a BDT fracture when applied stresses and resultant strains exceed certain limits, while rocks below deform in a plastic manner with little if any fracturing. Anomalously high temperatures or pressures at depth generally raise the BTD, which, for example, is much shallower than usual in "hot orogens" like magmatic arcs and backarc basins. Similarly, a cold slab of oceanic crust falling into the upper mantle drags the BTD down with it until it warms enough to lose all brittle behavior, or to ambient temperature, whichever occurs first.
The primary BDT within the basement defines the boundary between the upper and lower crust, but the cover may develop its own secondary BDT above the basement. Consider a typical monolithic Rocky Mountain basement block bounded by steep faults inherited from pre-Laramide tectonic events and now mantled by ~12,000' of flat-lying late Pennsylvanian to Mesozoic sedimentary cover accumulated during and since the rise of the Ancestral Rocky Mountains at ~300 Ma. The time is the latest Cretaceous, say ~70 Ma, and the block has just begun to rise along its reactivated (and now reverse) bounding faults at the onset of the Laramide orogeny. Due to the great thickness of the cover, high confining (lithostatic) pressures establish a secondary BDT somewhere above the basement-cover contact, while the primary basement BDT remains well below it. If a block margin above the basement BDT happens to fail as the block rides up and over adjacent blocks, it will break up in a brittle manner, just as the Rampart Range broke off the southeast margin of the Front Range. In contrast, the cover layers below the secondary BDT will drape over broken block margins like so much "melted cheese" on a chopped potato. Additional examples of ductile cover deformation over brittle basement fragmentation occur all along the east margin of the Front Range north of Boulder.
Acknowledgment: Thanks to Vince Edwards for introducing me to this compelling model of range-front deformation.
In addition to the home page references, this article relies on the following sources, in alphabetical order by first author:
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