Few natural landscapes on Earth match the Grand Canyon of the American Southwest for scenic beauty; and fewer yet rival its precipitous depths for the sheer magnitude of geological features on display or the profound nature of the story revealed.  Carved by the Colorado River through recently uplifted terrain, its sheer cliffs and ramparts offer an unprecedented opportunity to observe a spectacular sequence of sedimentary and crystalline basement rocks and geologic structures unparalleled among natural landscapes worldwide.  Any casual visitor to the Grand Canyon can see the magnificent array of rocks, lots of rocks, simply by looking over the canyon’s rim.  Igneous and metamorphic rocks are confined to the inner gorge of the canyon where the Proterozoic crystalline basement of the Granite Gorge Metamorphic Suite and Zoroaster Plutonic Complex have been deeply incised (Figure 1); but the bulk of the Grand Canyon’s exposures consist of sedimentary rocks of two distinct age groups.  The older, Late Proterozoic sedimentary rock sequence is comprised of isolated patches of Grand Canyon Supergroup, structural wedges of northeast-tilted rock layers bounded by northwest trending normal faults and unconformities at the base and top (Figure 1).  The younger, Paleozoic sedimentary rock sequence lies in layer-cake fashion above the Great Unconformity, superimposed on crystalline basement and tilted Supergroup rocks alike and spanning the majority of rocks expressed in base relief in the canyon’s walls (Figure 1).

The sedimentary rocks exposed in the walls of the canyon can tell us much about their origins if we know how to read the clues that they preserve. Sedimentary rocks are derived from the weathering and erosion of preexisting rock to produce sediment which is comprised of either the eroded “grains” of former rocks (referred to as clastic sediment), crystals precipitated from saltwater solutions (referred to as chemical sediment), or the organic detritus of decomposing plants (referred to as biological sediment); sediment that eventually becomes lithified to produce solid rock once more. Grains of clastic sediment can be described by textural features such as the size, shape, and orientation of the particles; textures that are used in their classification and to infer their maturity (degree of transport and recycling).  Chemical sediments are defined by their composition; the presence of fine to coarse crystals of carbonate, sulfate, halide, and other salts.  Carbonate rock classification also includes the relative abundance of allochems (visible fossils, ooids, pellets, and intraclasts). 

In addition to the compositional (or lithological) and textural properties of sediment, sedimentary rocks contain an array of sedimentary structures and fossils of plants and animals that can be used to reconstruct where, how, and when they formed.  Compositional and textural properties, structural elements, and fossil content are combined to describe a sedimentary facies, a unique suite of physical, chemical, and biological characteristics that a geologist can use to interpret the depositional environment in which the sediment (and thus, the sedimentary rock) accumulated.  Reconstructing a three-dimensional arrangement of facies (the vertical and lateral changes in facies characteristics) allows the geologist to properly infer the paleo-geographic and -climatic features of the (sedimentary) basin in which the sediment was deposited, as well as the tectonic controls over sediment source, type, production rate, and distribution; all independent facets of knowledge synthesized to provide a geologic history of a given region.

Figure 1.  The suite of rocks exposed by the downcutting of the Colorado River in Grand Canyon National Park.

Interpreting geologic history from sedimentary rocks requires an understanding for what happens to sedimentary particles as they are generated, as well as the processes by which sediments form.  What changes are wrought upon sedimentary particles broken down by weathering, abraded and further degraded by transport, accumulated in their “final” depositional environments, and eventually lithified back into solid rock?  What physical, chemical, and biological processes affect these changes over distance and time?  Fundamentally, it is important to understand that all of earth’s crustal materials, rocks and minerals, are continually undergoing decay, and the final products of those processes of degradation are always the same regardless of the complexity of the original material. 

It may help to begin with a “simple ideal model” for the evolution of sedimentary rocks (Figure 2), one that assumes that all processes which generate sediment and alter particles during weathering, transport, and deposition are eventually driven to completion (although multiple cycles of subaerial exposure, weathering, transport and deposition may ultimately be required) (Fichter and Poche, 2001).  Rock-forming minerals (those that crystallize or are recrystallized to form igneous rocks and many metamorphic rocks) are only stable (not prone to decay) under the conditions at which they form; a change in these initial conditions causes physical and chemical alteration to more stable products, the sediment of sedimentary rocks.  When igneous rocks form at the earth’s surface by volcanism, and when igneous, sedimentary, and volcanic rocks are exhumed at the surface by uplift (related to mountain building and isostatic adjustment of the crust), they degrade by weathering and erosion to remain as only the most stable of the preexisting minerals, or they form new, more stable minerals altogether.  When these rock-forming minerals undergo chemical decomposition, minerals that were crystallized first under higher temperatures and pressures, break down most rapidly because they exist at conditions most unlike those under which they formed.  Using Bowen’s Reaction Series to simulate the crystallization of rock-forming (Figure 9 from GEOLOGY BASICS), the weathering of all minerals will result in three basic end-member sedimentary components: quartz sand, clay, and ions in solution precipitated as calcite (Figure 2).  Quartz is the most stable at the earth’s surface, and is only slightly affected by weathering.  Quartz mineral fragments are freed from the rock matrix, broken down and rounded to become sand, a significant component of sandstones, as well as a lesser component of mudrocks.  The ferromagnesian minerals (minerals rich in iron, magnesium, and calcium – olivine, augite, hornblende, and biotite) decompose rapidly to form soluble salts (calcite) and small amounts of silica and clay.  Muscovite mica is relatively stable, and only small amounts are dissolved.  Muscovite fragments cleave and are broken into very small flakes; the smallest flakes sometimes undergo further change to become clay minerals.  Feldspars are generally decomposed rapidly to form clay minerals such as kaolinite and soluble cations of calcium, sodium, and potassium that are carried out to sea in solution where they recombine with other anions and may precipitate as new minerals such calcite, dolomite, or gypsum.

Figure 2.  The Simple Ideal Model for the evolution of sedimentary rocks (modified from Fichter and Poche, 2001).

The generation of sediment involves erosion and dissolution (the removal and transportation of sedimentary particles and chemical ions from their original host rock to a place of deposition by streams, wave action, tides, wind, and glaciers).  The three weathering products described by the simple ideal model (Figure 2) do not remain in a heterogeneous mix at their source; rather, they are transported down system (from land to the sea) and separated from each other through a process called sorting.  In our model, sand is relatively heavy and remains behind in higher energy environments such as the beach, while clay particles are light and remain in suspension even under fairly low energy (quiet water) conditions where it finally settles to the sea floor on the near shelf.  Chemical ions such as Ca+2 and CO3-2 travel in solution to the far shelf and remain there until high ionic concentrations allow chemical bonding and direct precipitation of calcite crystals from the water column.  More often, calcareous algae and marine animals extract the ions from sea water through the biomineralization process, and use them to build skeletons of calcite.  After death, skeletal material accumulates on the sea floor as carbonate sediment.

Earth’s sedimentary systems are messy and many transitional processes (including recycling of previously formed sediment and sedimentary rock) and gradational or transitory sediment types exist between the unweathered source rock and the three end member products of the simple ideal model.  Physical and chemical weathering act continuously to break down the original rocks and minerals and their semi-stable byproducts into the more stable end components.  Abrasion and breakage gradually removes the sharper, more angular edges of sedimentary particles and reduces their overall size.  Mechanical sorting of sediment during transport works to produce a regular hierarchy of materials under decreasing energy conditions (larger particles drop out first): ions in solution, clay and iron/aluminum oxides, sand grains, and larger lithic (rock) fragments.  Overall, as sediment is carried further down system, rounding and sphericity of particles increases, size decreases, and the homogeneity of size and composition increases.

Sedimentary Rock Classification

Sedimentary rocks are those that have formed as aggregates of clastic material eroded from pre-existing rocks, aggregates of bio-chemically and/or chemically precipitated minerals that accumulate in bodies of water or that form from the deposition of organic (mainly plant) detritus where decomposition is minimal.  Sediments are later lithified (buried, compacted, and cemented together) to become sedimentary rock.  A basic system of classification subdivides sedimentary rocks into three categories related to their formative processes and fundamental components: siliciclastic (comprised of particles derived from preexisting rocks), chemical and/or biochemical (comprised of crystals precipitated from salt water or biomineralized skeletal material, and biological (comprised mainly of plant remains).  Much concerning the basics of sedimentary rock identification and classification has been discussed elsewhere (see GEOLOGY BASICS) and need not be repeated here.  Instead, I devote this section to a refinement and expansion of the classification for siliciclastic sedimentary rocks and chemical/biochemical carbonate rocks.  These rocks are widely abundant and comprise the vast majority of the sedimentary rocks exposed in the Grand Canyon region.  

Clastic sedimentary rocks are classified on the basis of their texture (the size of the particles found within the rock) and their composition.  Textures or particle sizes are measured according to the Wentworth Scale (Figure 3), and are divided into three main classes: gravel (diameter > 2 mm), sand (diameter between 2 mm and 1/16 mm), and mud (diameter < 1/16 mm).    Each size category can be further subdivided; for example, mud consists of silt (diameter between < 1/16 mm and 1/256 mm) and clay (diameter < 1/256 mm).  The abundance of each particle size class must then be estimated; this can be determined with some practice using a relative abundance chart similar to the one shown in Figure 4.

Figure 3.  The Wentworth Scale of siliciclastic textures.

Figure 4.  A chart depicting the relative abundance of sedimentary particles.


Several classification schemes are available, each defining siliciclastic rocks in slightly different ways; however, several simple rules can be followed to make classification easier.  Geologists typically examine clastic rocks to determine concentrations of five compositional components: quartz, feldspars, lithic (rock) fragments, and mud matrix (silt and clay particles too small to identify their mineralogy in hand specimen).  Siliciclastic rocks containing at least 30% gravel (typically rock fragments) sized particles are called conglomerates (if the particles are dominantly angular, the rock is called breccia).  Conglomerates can be subdivided into those that are grain-supported (clasts are mainly touching) and matrix-supported (clasts are mostly surrounded by mud).  The classification of conglomerates and sandstones can be further refined using the relative abundance of quartz, feldspar, lithic fragments, and matrix.  Figure 5 shows a three-dimensional ternary diagram often used to classify these rocks (this is modified from Williams, Turner, and Gilbert (1982) to include both conglomerates and sandstones).  The front facing triangle has apices depicting the maximum abundance (100%) of quartz (top), feldspars (left base) and lithic fragments (right base); the third dimension to the triangle represents the abundance of matrix (mud sized particles).  Proper interpretation of the diagram takes some practice: conglomerates and sandstones containing 5% or less matrix are called arenites, but if they contain 5 to 50% matrix they are called wackes.  If the coarse fraction (gravel and sand) is 90% or more quartz, the rock becomes a quartz arenite or quartz wacke; if the coarse fraction has less than 90% quartz, but has greater feldspar content than lithics, the rock becomes a feldspathic arenite or feldspathic wacke; and if the coarse fraction is less than 90% quartz, but greater lithic content than feldspar, the rock becomes a lithic arenite or lithic wacke.  Incidentally, Figure 5 also defines a category of siliciclastic rocks called mudstones.  If there is more than 50% matrix (but also less than 30% gravel), the rock is neither a conglomerate nor a sandstone, instead, it becomes a mudstone.  Mudstones can be subdivided on the basis of silt versus clay content: if the mudstone contains 1/3 or less silt, it is called a claystone; if it contains 2/3 or more silt, it is called a siltstone; and if it contains between 1/3 and 2/3 silt, it remains simply a mudstone by name.  Silt is too small to see individual grains, but particles are large enough to impart a gritty feel to the siltstone; claystones feel smooth to the touch.  Mudstones exhibiting laminations (very thin, but visible layers within the rock) are usually referred to as shales. 

Figure 5.  Classification of siliciclastic rocks based on the four compositional components quartz, feldspars, lithics, and matrix (modified from Williams, Turner, and Gilbert, 1982).

Carbonates consist of a group of compositionally related rocks that share the same complex anion CO3-2.  Limestones are mainly formed of the mineral calcite (CaCO3) which is chemically bonded and precipitated directly from sea water or are biomineralized to form the skeletal structures of marine organisms.  Dolostones consist primarily of the mineral dolomite (CaMgCO3); Mg+2 ions readily substitute for Ca+2 ions in calcite crystals to form dolomite.  This substitution process is probably chiefly a diagenetic alteration or secondary mineralization that occurs after the deposition of primary calcite.  Regardless, dolostones generally are more prominently crystalline and are denser or more strongly indurated than are limestones, but the secondary mineralization process often does not completely destroy the primary features of the original limestone making classification possible.  Both limestones and dolostones typically contain a component of siliciclastic mud as well, although the mud does not factor into their classification.

The carbonate classification system I am most comfortable with was developed by Folk (1962), and is generally referred to as the Folk classification system for carbonate rocks.  Carbonate sedimentary rocks are classified on the basis of two characteristics: the allochems or the large pieces of fossils and other material they contain; and the interstitial material or matrix found between the allochems (Figure 6).  Two categories of matrix are utilized in the classification: micrite, which is a dull, grayish, lime mud comprised of clay and clay-sized crystals of calcite; and sparite, which is mainly a coarsely crystalline matrix of clear to translucent calcite minerals.   The initial rock name thus becomes micrite or sparite; although Folk’s classification does allow for carbonate rocks with mixed compositions which are called dismicrites.  Four categories of allochems are then identified and used to modify the rock name: 1) fossils that may be whole, broken, and possibly abraded; 2) oolites consisting of small, spheriodal, concentrically layered grains of lime mud; 3) pellets formed of the fecal matter of invertebrate marine organisms; and 4) intraclasts, pebbly chunks of indurated lime mud eroded and redeposited by wave action.  The prefix “bio” is added to micrite or sparite; if fossils are present, as well as the prefix “oo” for oolites if present, the prefix “pel” for pellets if present, and the prefix “intra” for intraclasts if present.  If two or more allochems are present, the allochem of greatest abundance is placed closest to the name micrite or sparite.  Thus, if the carbonate rock exhibited a sparry matrix and contained abundant fossil fragments with a minor component of pellets, it would be named a pelbiosparite. 


Figure 6.  Folk’s carbonate rock classification system (after Folk, 1962).

Two other features can be used to further characterize the carbonate rock; relative abundance and size of the allochems.  The relative abundance of allochems can be indicated by describing the rock as either grain-supported (allochems are touching) or matrix-supported (allochems are floating in a fine interstitial material).  Allochem size is based on the Wentworth Scale, although the term rudite is substituted for gravel-sized allochems, calcarenite for sand, and lutite for mud (silt and clay).  These names are added as a suffix (substituting “ite” in the initial name of the carbonate rock.  In this way, the name of a carbonate rock exhibiting abundant, coarse, fossil allochems with a minor component of pellets in which the majority of allochems are in contact would be described as a grain-supported pelbiosparudite; wow, that’s a mouthful!   Finally, a separate category of carbonate rock containing reef-forming organisms such as colonial corals or stromatolites are termed biolithites under Folk’s classification (Figure 6).  No indication of abundance or size is suggested, although the rock must exhibit the framework nature of the organism to be identified as a biolithite.

Other Descriptive Features of Sedimentary Rocks

Sedimentary rocks contain many other features significant to reconstructing their depositional environment.  Describing the textures and composition of siliciclastic rocks is useful for interpreting the degree to which they have been transported and/or recycled; and thus, their progress toward achieving the siliciclastic end members indicated by the simple ideal model (Figure 2).  These alterations to the sediment as it travels down system and through time are known as maturity.  The maturity of sedimentary particles is measured by changes in the size, shape, and sorting of grains due to abrasion and breakage (textural changes) and changes in the content of lithic fragments and minerals due to weathering (compositional changes).  Textural maturity is measured against the ideal end products quartz sand and clay; greater maturity is indicated as grains become smaller (Figure 3), more rounded and more spherical in their overall dimensions (Figure 7), and better sorted (Figure 8).  A texturally mature siliciclastic sediment exhibits grains that have been worn down to nearly perfect spheres and sorted to nearly uniform size.  A compositionally mature sediment is one in which the lithic fragments and unstable minerals have been reduced to their ultimate end member compositions, quartz sand and clay mud. 

Figure 7.  Grain shape; sedimentary particles become more rounded and spherical as they are transported.

Figure 8.  Grain sorting; sedimentary particles become more uniform in size, shape, and composition as there are transported.

Sedimentary rocks of all types (not just siliciclastics and carbonates) often exhibit sedimentary structures, particular patterns or arrangements of sedimentary particles created by certain physical, chemical, and/or biological processes that occur during deposition or subsequent to deposition.  Primary structures are formed during sedimentation and are related to the flow regime of the environment that the sediment was deposited in.  Structures that form after deposition, but before the sediment is lithified, structures formed by disturbance of previously deposited sediment, are known as penacontemporaneous structures; and structures that form after lithification that may have little to do with the original environment of deposition, including concretions, color banding, and weathering rinds, are known as secondary structures. Primary and penacontemporaneous structures are generated by four mechanisms generally related to the depositional setting where they formed: 1) depositional structures, formed mainly during deposition; 2) erosional structures that involve an episode of erosion followed by deposition; 3) deformational structures that involve deposition, followed by physical disturbance of soft, water saturated sediment; and 4) biogenic structures formed by the modification of sediment from biological organisms.

Primary sedimentary structures occur in response to flow regime which represents the range of energy conditions that transport sediment.  Transport occurs in certain ways under specific energy conditions and results in distinctive sedimentary structures that characterized that mode of transport.  Flow regime is a continuum between two end members, between unidirectional flow, which is flow in one direction (like a river), and oscillatory flow, which is flow that alternates direction (like waves on a beach).  Although these end-members are common enough under natural conditions, more often than not, they occur in tandem under combined flow conditions.  Primary structures include bedding (or stratification), which describes the way in which sediments accumulate in layers, as well as bed forms, which occur at the interface between individual beds or layers.

All sedimentation processes arrange sedimentary particles into three-dimensional “packages” called beds that can be distinguished from layers above or below by the composition, texture, or color of the sediment.  These structures can be described by their external morphology which includes the size, shape and orientation of the bed relative to other beds within a particular stratigraphic unit (a body of rock with similar characteristics), and internal morphology, which includes features within the bed’s interior.  Bedding thickness is variable, beds can be thin (between 1cm and 10cm), medium (up to 30cm), thick (up to 100cm), and very thick (greater than 100cm).  Beds thinner than 1 cm are referred to as laminae or laminations, and these are described as very thin (up to 0.1 cm), thin (up to 0.5 cm), and thick (up to 1.0 cm).  External shape and orientation of the body of sediment can be described as tabular, wedge, lensoidal, or irregular.  Internal morphology describes laminations contained within a bed and includes planar (flat-lying), cross (tilted), wavy, and convolute (irregularly swirled) laminae.  A bed with cross laminae can exhibit laminations that are planar (angular or tilted at an angle relative to the bed), concave (tangential to the bed), trough (dish-shaped), or hummocky (gently undulating swales).  Cross-bedding forms by the migration of ripples/dunes in water or air (Figure 9).  Hummocky cross-bedding, characterized by undulating sets of cross-laminae alternating between concave-up and convex-up cut gently into each other to form curved bounding surfaces.  These features form by oscillatory or combined flow associated with large storm surges that only occasionally affect deeper water settings.  A bed is named by first indicating the thickness, then external morphology, and then internal morphology.  For example, if preserved in the rock record, the beds depicted in Figure 9 could be described as “very thick, tabular sets of tangential cross laminae.” 

Figure 9.  Crossbedding produced by dune or ripple migration; in this case described as very thick, tabular sets of tangential cross laminae.

Groups of vertically stacked beds of similar composition, internal structure, etc. are called bedsets and are usually genetically related (often representing a sequence of material deposited in associated environments).  The boundaries (bedding planes) between beds are usually marked by a change in texture or composition and may be sharp or gradual.  Bedding planes represent periods of erosion, nondeposition, and/or changes in depositional conditions; erosional surfaces are obviously more abrupt.

Parallel bedding refers to beds that do not contain internally dipping laminae (cross-laminae) that are bounded by nearly flat, parallel bedding planes. These beds may contain internally parallel laminae, particle size grading, or they may be massive (with no internal structure).  Their origin is often linked  to deposition from suspension, changes in external conditions (such as climate or sea level), or changes in the energy conditions associated with deposition.  Parallel beds that exhibit grading of particles are referred to as graded bedding, characterized by sharp basal contacts and either an upward decrease in grain size called “normal” (Figure 10), or an upward increase in grain size called “reverse”.  Normally graded bedding is fairly common and usually associated with changes in the density of the material as it is deposited from suspension, whereas reversely graded bedding is rare and often related to pyroclastic flows, mass-wasted grain flows, and density segregation of fine grained heavy minerals from coarse grained light minerals in beach environments.

Figure 10.  Normally graded bedding exhibits a decreasing grain size upward through the bed, and is formed during the waning of a current and/or by settling of suspended particles of different size and density under low energy conditions.

In addition to bedding, several prominent bed forms can develop along bedding plane boundaries (recall that internal cross-laminae form by migration of different bed forms).  Ripples (small bedforms) and dunes (larger bedforms) occur in both siliciclastic and carbonate sediments and form by both water and wind transport of granular particles.  Ripples and dunes develop in either unidirectional or oscillatory currents (Figure 11).  When formed in unidirectional flow, they are asymmetric in shape, with the steep (lee) face of the ripple or dune crest in the down-current direction.  Unidirectional currents are related to stream flow, beach backwash, longshore drift, tidal currents, and deep-ocean bottom currents.  Oscillation ripples form by wave action and tend to be symmetric in shape with straight crests.  Ripple and dune migration leads to formation of dipping foreset laminae and bottomset laminae (more tangential to the bed) by avalanching and suspension settling in the zone of separation on the lee side of the bedform (Figure 9).  Foreset laminae are generally inclined more steeply than bottomset laminae (especially when foreset laminae form from avalanching of coarser particles down the lee slope onto the bottomset laminae).  The preservation of crossbedding is much greater than the bedforms themselves because the migration process involves planning-off the upper portion of previously formed ripples and dunes.

Figure 11.  Asymmetric and symmetric ripple marks associated with unidirectional and oscillatory currents, respectively.


Erosional structures may form within the upper portion of a previously deposited bed.  Channels having U- or V-shaped cross-sections that cut across previously deposited material (and other structures) are generally preserved by in-filling with other sediment.  They may form by fluvial and tidal currents, or by submarine gravity flows and are generally longer than they are wide and subparallel to flow direction.  Scour-and-fill structures are smaller, asymmetric-shaped troughs with long axes oriented parallel to flow and a steeper up-current slope than down-current that form by initial current scour and subsequent backfilling as current velocity decreases or shifts.  In contrast, sole markings are positive relief features formed in a two-step process on the base of beds.   First, a layer of cohesive, fine-grained sediment is eroded by some mechanism to produce elongate grooves and/or depressions; and second, a coarser-grained sediment fills the groove/depression to preserve it.  These features may result from current scour or from the action of tools (the current induced drag or roll of pebbles, pieces of wood, or other material along the bottom).

Recall that penecontemporaneous structures form by disturbance of previously deposited sediment before the sediment has lithified.  Many of these structures have been identified, but they generally fall into two groups; those related to the physical conditions within the depositional environment, and those related to biological conditions (bioturbation of the sediment after deposition).  The former category includes a diverse set of structures too numerous to describe, but several types of loading and dewatering structures, as well as features such as mud cracks and raindrop imprints occur.  Loading of fresh sediment onto partially liquefied sediment soon after deposition can generate convolute bedding and laminations, the intricate folding and crumpling of sedimentary layers into small, irregular, anticlines and synclines.  Similarly, highly irregular swellings or knobby protuberances on the sole of sandstone beds overlying mudstones, referred to as load casts, can form as less dense mud is squeezed up into the sand.  More intense loading creates flame structures (Figure 12), wavy or flame-shaped tongues of mud that project upward into an overlying layer of coarser material.  On the other hand, dish and pillar structures are generated in association with compaction and dewatering of sediment.  The “dish” is a thin, dark, flat to concave-up, clayey lamination in sandstones and siltstones; and the “pillar” is a vertical, cross-cutting column of structureless or swirled sand.   Subaerial exposure allows desiccation of mud and muddy carbonates to form upward tapering, V-shaped, crudely polygonal fracture patterns (Figure 13). Subsequent infilling of the cracks with fresh sediment forms positive-relief fillings on the tops of bedding surfaces called mud cracks (although they are actually the mold of the crack, not the crack itself). Exposure of soft sediment to raindrop or hailstone impacts can also generate small, crater-like pits called impressions on bedding plane surfaces.  Organisms may disturb soft sediment after deposition (animals mainly, but plants as well).  These bioturbation features are known as trace fossils (Figure 14) and include the tracks, trails, and burrows of animals, as well as root casts, bark impressions, etc., left by plants.

Figure 12.  Loading casts in the Tapeats Sandstone, generated by the accumulation of a denser sediment (sand) above saturated, less dense, previously deposited material (mud).

Figure 13.  Polygonal fracture patterns formed by the desiccation of mud and subsequent infilling by fresh sediment are called mud cracks; these are preserved in the Hakatai Shale of the Grand Canyon Super Group.

Figure 14.  Trace fossils; in this case, fossilized trackways of a small desert-dwelling reptile preserved in the Coconino Sandstone (A), and marine worm borrows preserved in the Bright Angel Shale (B).

Bodies of sedimentary rock can be distinguished on the basis of their lithology (texture, composition, and sedimentary structures) and organized into lithostratigraphic units; stratified, tabular bodies bounded by distinct (sharp/abrupt) or gradational contacts.  Formations are the fundamental lithostratigraphic unit and may consist of singular or multiple interbedded, intertonguing lithosomes; where a lithosome is a body of rock with essentially uniform character, distinguishable from other bodies above, below, or adjacent, that is large enough in scale to be mappable (thick and widely distributed).  Formations can be subdivided into members and/or beds; and joined into groups and/or supergroups; thus forming a regular hierarchy of increasingly complex, thicker, and more widely distributed rock bodies from bed to member, member to formation, formation to group, and group to supergroup.  Recall that Figure 1 shows the (lithostratigraphic) rock units exposed within the confines of the Grand Canyon.

Lithostratigraphic units are separated from each other by vertical contacts (planar or irregular surfaces between layers of different types of rock).  Conformable contacts occur where strata are superposed in an unbroken depositional sequence (the contact is a conformity, a surface with no evidence of nondeposition or erosion).  Unconformable contacts occur where strata are separated by distinct breaks in deposition called a hiatus, which may include erosion (the contact is an unconformity).  Unconformities are covered in the section on GEOLOGY BASICS.  Vertically conformable contacts may be abrupt; separate beds of distinctly different lithology caused by minor changes in depositional conditions and/or postdepositional chemical alteration.  More likely, they are gradational, exhibiting less marked changes in lithology.  Progressive gradual contacts occur where one Lithology grades into another by gradual changes in physical characteristics or composition, while intercalated contacts occur where an increasing number of thin interbeds of another lithology appear upsection.  Lithostratigraphic units do not extend indefinitely; they terminate abruptly by erosion or more gradually by change in depositional environments.  Lateral pinch-outs occur where lithologic changes exist as progressive thinning of a unit to extinction.  Intertonguing occurs where lateral changes exist by splitting of a unit into many thin units that pinch out independently.  Progressive lateral gradation occurs where lateral changes gradually shift from one lithology to another.

Depositional Environments

Sediments and sedimentary structures rarely occur as isolated features in sedimentary rocks, if so, they would not be particularly useful for inferring (reconstructing) the conditions of the depositional environment from whence they came.  Instead, features of sediment maturity (texture and composition), and sedimentary structures often occur in genetically (spatially and temporally) related, regularly repeated, vertical successions called stratigraphic sequences that represent fluctuating flow-regimes over time as energy conditions dissipate or shift position.  It is important to understand that geologists have distinguished, modeled, classified, and can now recognize virtually every depositional environment on the basis of its characteristic sediments and sedimentary structures.  That is, specific sedimentary sequences are associated with specific depositional environments; when a sedimentary sequence is identified, its consequent depositional environment is identified.  Geologists use the concept of sedimentary facies to describe the sum total physical, chemical, and biological characteristics imparted to a sedimentary rock at the time of its deposition.  But facies has another meaning as well, it is a body of sedimentary rock (a single bed or multiple beds) with a distinct suite of characteristics that was generated in a depositional environment; thus, facies are the many different sediments and resulting rocks that form at the same time in different depositional environments.  Returning to the Simple Ideal Model (Figure 2) of Fichter and Poche (2001), the quartz sandstone, clay mudstone, and calcite-rich limestone are the facies formed in beach, near shelf, and far shelf depositional environments. 

The context of any facies, that is, its relationship to other facies, must be known before an environmental interpretation is made.  These relationships are called facies associations, and in general, any depositional environment is characterized by multiple facies (i.e., the facies association) that when grouped together form a “model” describing that environment.  A series of facies which pass vertically from one to another gradually is called a facies sequence or succession (correlative with the stratigraphic sequence discussed earlier).  Most successions are bounded at the base and top by a sharp or erosive truncation or by a depositional hiatus; therefore, facies occurring in a sequence conformably above one another (in gradational contact) must have formed in laterally adjacent environments.  This is an important concept in stratigraphy known as Walther’s Law of Facies.  Gradational contacts indicate the migration of adjacent depositional environments; on the other hand, sharp or erosive contacts indicate facies that may have formed in environments that were widely separated in space or time.  Vertical sequences may not only reflect changes brought about by natural sedimentary processes internal to a basin, but they may also reflect changing external controls.  Interbasinal controls may include progradation of a delta, the lateral migration and avulsion of a stream channel, or the buildup of a carbonate platform/reef complex; whereas extrabasinal controls could involve changing sediment supply, sea-level, climate, and/or tectonic activity.  Using the Simple Ideal Model to represent facies associations on a continental shelf margin, Figure 15 shows the shifting positions of laterally adjacent depositional environments produced by changing sea level (marine transgression and regression).  Taking a vertical slice through the sedimentary succession at various locations indicates how migration of laterally adjacent facies generates vertical onlapping and offlapping of facies sequences.  Onlapping sedimentary sequences produced by transgression (Figure 15a-c) are also know as deepening upward or fining upward sequences, whereas offlapping sedimentary sequences produced by regression (Figure 15d) are known as shallowing-upward or coarsening-upward sequences.  The latter sequence can also result from progradation of environments into a basin with stationary sea level.

Figure 15.  Facies associations and sequences of the Simple Ideal Model generated by marine transgression and regression; a fining upward, onlapping facies sequence is produced by transgression (A, B, and C), while a coarsening upward, offlapping facies sequence is produced by regression or progradation (D).

It should be apparent from the preceding discussion of sedimentary rocks that sediments come in a much wider variety than the quartz sand, clay, and carbonate material indicated by the Simple Ideal Model (Figure 2).  Fitcher and Poche (2001) present two models expressing the diversity of sediments and the interconnectedness of sedimentary depositional environments which they term “long system” and “short system” models for the evolution of sedimentary rocks.  Both models portray the systematic changes in sedimentary rocks downstream from their sourceland of origin to their ultimate resting place in the ocean.  The models are idealized; they assume that sediments, sedimentary structures and sequences, and depositional environments evolve simultaneously, the sourcelands are compositionally and tectonically simple, and that the temporary storage and recycling of sediment within the system occurs in ideal depositional environments without significant transitions between them, from more proximal to the sourceland, to increasingly distal from the sourceland.


The Long System and Short System models are shown in Figure 16.  In the long system (Figure 16a), the sourceland is an uplifted mountain range produced by normal faulting or thrust faulting and formed in a continental interior consisting of typical, continental crystalline basement comprised chiefly of granitic igneous rock.  Sediments are ultimately carried from this sourceland to a distant depositional basin associated with an epicontinental sea lapping onto a passive continental margin.  The highland source primarily undergoes physical weathering and mass wasting to form coarse, angular debris deposited at the mountain front by braided distributary streams and debris flows to form alluvial fans where stream gradient and transport energy decrease rapidly (Figure 17). Alluvial fan facies comprise a complex integrated mix of fluvial to debris flow dominated environments that generally accumulate immature, matrix-supported, feldspathic breccias.  Some of the material remains to form part of the geologic record as sedimentary rock, but much of the sediment is recycled and transported down system.  During transport the material is broken up, abraded, and rounded to become immature, grain-supported, feldspathic conglomerates deposited as proximal, braided stream facies.  The braided river facies sequence consists of alternating channel L-bars (longitudinal bars) comprised of gravels and T-bars (transverse bars) comprised of cross-bedded sands (Figure 18).  Braided rivers have no true flood plain; channel deposits typically occupy the entire valley floor over time.  Sediment transported during higher flows is deposited as water level and energy decline; L-bars form within the channel center parallel to flow first, forcing the channel to split or braid, whereas T-bars form perpendicular to flow during further waning of energy conditions.  L-bars are gradually replaced by T-bars in proximal to distal braided rivers.  

Figure 16.  Long system (A) and short system (B) models for the evolution of sedimentary rocks (modified from Fitcher and Poche, 2001).

Figure 17.  Facies model for a typical proximal system, debris flow-dominated alluvial fan.

Figure 18.  The typical proximal system, braided river facies sequence.

In alluvial fan and proximal braided stream environments where gradients are steep, energy conditions are high with sediment moving by a combination of gravity-induced sliding and rolling, and flowing water.  Further down system, where gradient drops off, there is less gravity-induced transport of sediment and greater force of running water.  The flow of water exerts greater force during heavy rain and stream flooding to move coarser material, but during normal flow conditions, gravel is sorted out and left behind upstream.  Sediment is transformed to become immature, feldspathic arenite sandstones and deposited as distal, braided stream facies.

Chemical weathering begins to replace physical weathering as a major process down system.  Feldspars are decomposing to clay and ions in solution, material that is easily carried by running water even under gradually decreasing stream gradients and energy conditions.  Feldspathic arenites are altered to submature, feldspathic wacke sandstones and deposited as meandering stream facies as the amount of feldspar grains declines and the relative proportion of quartz grains increases.  The meandering river facies sequence consists of fining upward layers of channel, point bar, and flood plain deposits (Figure 19).  Each bed in the sequence is deposited under steadily decreasing energy as individual loops in the meandering river migrate.  The channel undercuts the bank comprised of flood plain material opposite the point bar, forming a lag of mud pebbles (channel gravels) on an erosional base.  Channel migration ensues.  Initial loss of energy allows coarse to medium sands to accumulate as trough cross-beds and high energy planar laminations.  Point bar deposition begins with medium to fine trough cross-beds with ripple marked bed planes and ends with fine sand and silt containing common climbing ripple cross-laminae.  After the channel has swept beyond the area, flood plain muds accumulate, interspersed with root casts and mudcracks, and occasionally interbedded with fine, planar and trough cross-bedded sands (sands deposited by the higher energy waxing flood conditions and muds by lower energy waning flood conditions).  The sequence shown in Figure 19 would form during a single pass of a meander loop, sequences repeat in whole or in part as successive meanders migrate past.

Figure 19.  The typical medial system, meandering river facies sequence.

Continued weathering and abrasion during meandering river transport reduces sediment to the consistency of silt and clay particles mixed with quartz sand, resulting in a quartz wacke, although rivers carry a significant load of ions in solution as well.  Eventually, meandering rivers reach transitional deltaic and estuarine environments along the coast of an epicontinental sea.  Deltas form a complex of integrated depositional environments that include terrestrial distributary channel, crevasse splay, delta floodplain, and distributary mouth bar, as well as marine interdistributary bay, delta front, and prodelta facies difficult to represent in a single diagram (Figure 20).  Deltas are generally subdivided into three end-member types on the basis of fluvial, tidal, and wave dominance.  Figure 20a shows an aerial view of a delta influenced mainly by fluvial processes (as suggested by our Long System model).  As deltas evolve, they undergo a phase of active progradation, followed by an abandonment phase.  During progradation, mud is carried to the delta front and prodelta, while bars of sandy sediment accumulate at the mouth of distributary channels.  Distributary mouth bars build until they choke the original channel, forcing it to divide and eventually migrate beyond the bar (often cutting into it), only to accumulate new distributary mouth bars down delta.  Floods may cause temporary repositioning of flood water and deposition of sandy crevasse splays and finer overbank muds between distributary channels by breaching natural levee systems.  Tidal currents or wave action and nearshore currents can redistribute and further separate sediment along the delta front to form tidal flats or beaches.  During abandonment, the main distributary channel switches to a new location, shutting off the supply of sediment; the “old” delta gradually subsides below sea level as coal swamps inundate the area.  Figure 20b provides a stratigraphic section of a typical coarsening upward, prodelta to distributary mouth bar facies sequence associated with progradation of the delta into the marine basin.  

Figure 20.  A typical delta influenced by fluvial and wave processes; (A) provides an aerial view, (B) describes an idealized facies sequence produced by delta progradation.

In the Long System model, the high energy of wave action picks up and sorts the sand from the mud in the sediment brought to the delta to produce two types of mature clastic sediments; quartz sand remains behind to form quartz arenite sandstones deposited as beach and near shore facies, while silts and clays are transferred offshore to form mudstones deposited as near shelf facies of the passive continental margin.  The dissolved ions are brought to the ocean, diluted into the saltwater, eventually to be precipitated from the seawater as carbonate rocks, the far shelf facies of the passive continental margin.  Thus, the endmembers of the Simple Ideal Model (Figure 2), quartz arenites, mudstones, and carbonates are formed and deposited in their respective coastal, proximal shelf, and distal shelf environments.

Figure 28 in GEOLOGY BASICS provides a reasonable model of the beach and near shore environments of a typical coastline.  Assuming a stable sea level, gradual progradation of these environments into the basin would generate a coarsening upward facies sequence consisting of offshore, lower shoreface, upper shoreface, and foreshore units (Figure 21).  The offshore facies represents the transition to the near shelf environment; it lies below the normal (fair weather) wave base, but within the zone disturbed by storm waves.  Muddy sediment is deposited during pervasive quite water conditions, but storm waves agitate mud and transports in coarser silt and sandy sediment.  The lower shoreface environment generally lies within the coastal zone disturbed by fair-weather waves between high and low tide.  Sediment is commonly disturbed by wave activity; muds are usually sent seaward and low angle crossbedded sands accumulate within the zone.  The upper shoreface is routinely disturbed by wave action as it lies continuously above the fair-weather wave base.  Muds are willowed from sediment and transferred seaward while only trough-crossbedded sands remain behind. Seaward dipping planar beds of sand accumulate as foreshore facies where constant wave action in the surf zone constantly washes forward and back.  The facies sequence can be repeated multiple times in a steadily subsiding basin or when marine transgression is episodic. 

Figure 21.  A typical coarsening upward facies sequence produced by basinal progradation of distal system beach and near shore depositional environments.

The near shelf environment is dominated by one of two end-member transport processes, storm waves and tidal currents.  Where tidal activity is minimal, but a broad area of shelf lies between storm and fair-weather wave base, hummocky facies sequences accumulate (Figure 22).  In general, intermittent storms pass over the shelf, generating high energy wave conditions that disturb muddy sediment accumulated on the shelf floor and transport sandy sediment from the coast.  The bottom sediments are scoured and a basal layer of lag gravels consisting of indurated mud pebbles and/or shell fragments is deposited.  Gradually diminishing energy allows the deposition of hummocky cross-stratified sands, planar laminated sands, and sandy layers exhibiting small-scale trough cross beds (hummocky cross laminae are indicative of oscillatory currents created by wave action).  A return to quiescence between storms causes a new layer of mud to accumulate, often exhibiting bioturbation.  Where the near shelf environment is disturbed by strong tides, daily currents generally rotate in a circular path around the basin.  The currents generate and transport tremendous quantities of sand as large sheet-like bodies of various shapes and thicknesses.  Tidal shelf deposits often range from multiple stacked layers of moderately thick, tabular cross-bedded sand layers to thick, tabular mega-crossbedded sand bodies.  In either case, thinner, wavy and lenticular sand beds form at the sides and front of the moving sand body, while muddy sediment lies in between the migrating sands.  Figure 23 displays an idealized facies sequence formed by passage of a sand body. Initially, sandy patches move over muddy sediment to form basal lenticular sands.  More contiguous rippled sand sheets follow, depositing wavy sand beds.  Small sand waves, and then large sand waves cross the area, which are in turn partially eroded at the up-current end of the sand body’s migration to produce a capping gravel lag exhibiting basal flutes and scours.

Figure 22.  A model of the typical environmental setting and hummocky facies sequences associated with storm-dominated shelves (modified from Fichter and Poche, 2001).

Figure 23.  A model depicting migration of a tidal sand sheet and its associated facies sequence (modified from Fichter and Poche, 2001).

Alternatively, sediment generated from weathering and erosion of particularly quartz-rich bedrock localities, such as quartz pegmatites (a very coarse felsic igneous rock) or rock rich in quartz veins(precipitated from hydrothermal solutions), forms quartz gravels.  These gravels do not degrade readily and simply remain as lag deposits in the vicinity of their source until the mountains are worn down and reduced to a penaplain.  The sediment generally begins as feldspathic conglomerate and it gradually transformed to a wacke quartz conglomerate, and finally to pure quartz conglomerate (Figure 16a).     

Since continental crust is often a very complex amalgam of rock types and tectonic settings, it is quite possible the ideal long system presented above could start instead from a more complex sourceland of igneous, sedimentary, and metamorphic rocks associated with volcanic arcs and collisional mountain belts formed on an active continental margin (Figure 16b).   Physical weathering and mass wasting of this great complexity of rock types would tend to produce an abundance of lithic fragments (although weathering would still produce lesser amounts of feldspar grains too).  Systems draining the interior side of these mountains toward a passive marginal setting like the one described previously would consist of the same general environments linked in sequence: alluvial fans, braided rivers, meandering rivers, and epicontinental seas.  The abundance of lithic material would result in the production of lithic breccias, lithic conglomerates, lithic arenites, and lithic wackes (rather than feldspathic rock types).  Eventually, abrasion during stream transport and chemical weathering would reduce all lithics to individual mineral grains, and those too would be transformed to the quartz sand and mud of meandering rivers to form quartz wackes.  From this juncture onward, both long system paths merge into one.

Short systems drain rapidly from tectonically active mountain belts built adjacent to the deep, oceanic-trench-related depositional basins (Figure 16b), the exterior, volcanic arc side of the active continental margin described above.  Short systems can also form in association with volcanic island arcs developed on oceanic crust and fore-arc basins of collisional mountain ranges.  In these settings, the distance and time that sediment is transported is significantly telescoped, the number and length of depositional environments along the way is reduced, and therefore, the systematic changes in textural and compositional maturity related to the long system cannot fully develop.  The mechanical energy generated by gravity and flowing water is dissipated more quickly, making abrasion and sorting processes less effective.  Chemical weathering has less opportunity to break down unstable minerals.  As a result, sediments tend to be less mature; the material is coarser and less well sorted, and it tends to retain more feldspars and lithic fragments with a lower proportion of quartz grains.  As with the long system, the initial depositional environment of the short system is the alluvial fans formed at the base of the mountains, where deposition of immature, coarse-grained, angular, poorly sorted, matrix-supported lithic (or possibly feldspathic) breccias occurs.  Abrasion from downstream transport increases the rounding of clasts to produce lithic (or feldspathic) conglomerates and lithic (or feldspathic) arenites associated with proximal to distal braided river facies. 

Conglomerates and sandstones of the braided rivers are debouched directly into the transitional marine environments of the coastline.  Beach and nearshore facies here consist of lithic sandy conglomerates and conglomeratic sandstones (arenites).  Only a short distance offshore, the ocean floor drops rapidly along a steep slope to the deep basin many thousands of feet below.  Much of the lithic- and/or feldspar-rich sediment bypasses the narrow coastal zone with its limited accommodation space transported by turbidity currents to pour into the deep ocean, forming submarine fan facies called Bouma sequences composed of lithic conglomerates and lithic arenites in the proximal fan to lithic wackes in the distal fan (Figure 24).  Dark mudstones accumulate on the oxygen-poor ocean floor environment, beyond the outer margins of the submarine fans.  Bouma sequences are typical of submarine fans, although not exclusive to them.  Complete sequences consisting of the submarine fan facies ABCDE (Figure 24) form in mid-fan settings, while incomplete AE sequences are common to proximal channels, CDE or BCE sequences form adjacent to channels, and DE sequences form in more distal fan settings. 

Figure 24.  A model depicting a typical distal system submarine fan environment and its associated ideal Bouma sequence.

As indicated by the Simple Ideal model (Figure 2), carbonates can be generated distally in long siliciclastic systems (more rarely in short systems).  However, under these conditions, the volume of siliciclastic sediment usually overwhelms carbonate production, especially in coastal areas with high fluvial (and clastic) input, as well as short systems associated with active tectonics. Carbonate-dominated systems develop on passive continental margins where siliciclastic sourcelands are subdued and/or very distant, the water is generally shallow, and climates are tropical to subtropical.  In the warm, clear, shallow water, calcareous algal production flourishes, generating copious volumes of micritic (lime) mud, while invertebrate animal skeletons accumulate as fossil sedimentary particles.  Oolitic, pelloidal, and intraclastic allochems can also be produced in localized settings, depending on conditions.

Transitional and marine environments predominate in carbonate systems.  Figure 25, one possible model of carbonate environments described by Fichter and Poche (2001), presents a simplified system comprised of five “belts” running more or less parallel to the coastline: tidal flat, lagoon, barrier reef, shelf, and basin.  Tidal flat environments consist of supertidal (above normal high tide) and intertidal (between high and low tide) settings.  Supertidal facies most commonly occur as massive, dolomitized micrite deposits with lesser evaporites such as gypsum.  Micritic muds often exhibit mudcracks and algal laminates (stromatolite mounds).  Micritic muds with stromatolites are also common to intertidal facies, although carbonate sands exhibiting cross- and wavy- to lenticular-bedding form in shoaling environments.  Mud flats are crossed by erosional tidal channels containing flat-pebble conglomerates.  Lagoon environments lie within the coastal subtidal zone and contain a wide variety of interbedded sediments.  Micrites and biomicrites are commonly deposited in the calm, shallow waters, although a patchy distribution of storm deposited calcareous sand and coquina shoals washed into the lagoon from barrier reefs are fairly frequent.  Oomicritic and oosparitic sands formed from tidal sand bars and patch reefs are evidenced by the presence of mounded biolithic carbonates.  Calcareous layers may be interbedded with fine siliciclastics. The barrier reef protects the coast from the brunt of storm wave activity and its continuously agitated conditions are favored by reef-dwelling corals and other invertebrates that accumulate as biolithic limestones and massive, structureless biomic- and biosparites.  Subaqueous gravity flows, tubidity currents, debris flows, and slumps cascade down the front slope of the reef system to become deposited as bioclastic conglomerates.  The shelf environment may include shallow and deep shelves.  The shallow shelf, or carbonate platform is comprised of patch reefs, and storm-wave agitated muddy or sandy carbonate substrates.  Micrites are common, but are often interbedded with hummocky sequences of bio- and pelmicrites, thick layers of crossbedded bio- and oosparites or megarippled biosparites and pelmicrites.  A diverse suite of very abundant, whole fossils is often preserved in these deposits; and nodular cherts are not uncommon.  The rich complexity of calcareous deposits on the shallow shelf is replaced by substantial accumulations of interbedded, massive, bioturbated mudstone and micritic limestone on the deep shelf.  Deep shelf settings may contain carbonate-dominated submarine fan deposits transported in from adjacent barrier reef and platform environments.  The basin environment suffers from lack of light and oxygen and is typically not biologically productive.  Sediment accumulates as massive to thinly laminated dark (organic-rich) mudstones and micrites with rare silty or fine sandy distal submarine fan deposits.  Bedded cherts may form, especially below the carbonate compensation depth (the depth below which carbonates tend to dissolve back into cold, CO2 under-saturated ocean water).  Mudstones and cherts become more common in deep basins where they are unmasked by the limited deposition of carbonates.    

Figure 25.  A simplified model of a carbonate-dominated depositional system consisting of five proximal to distal environmental belts: tidal flat, lagoon, barrier reef, shelf, and basin (modified from Fichter and Poche, 2001).    

Sedimentary Tectonics

Any study of sedimentary rocks would be incomplete without an analysis and interpretation of their relationship to plate tectonics and the clues to geologic history that they hold.  As indicated previously in my discussion of long and short siliciclastic systems, sedimentary rocks are not random collections of earthy material, lodged willy-nilly about the earth’s surface; but instead, the characteristics they preserve are inherited from their sourcelands and from the depositional basins where they accumulate.  Tectonic formational processes controls in large part both the type and composition of the crustal uplift from which sediment is generated by weathering and erosion, and the parameters of the basin where the sediment accumulates to form sedimentary rock.  Each type of sedimentary rock outcropping in a depositional basin preserves compositional and textural features, sedimentary structures, and fossil content that reflects the physical and geological signature of its sourceland (relief, composition, distance from the basin, etc.), as well as information about the unique environmental conditions of the basin of deposition itself, it is simply up to the geologist to tease out a geologic history from the evidence which remains.

A complete description of the intricacies of sedimentary tectonics is beyond the scope of this discussion, but broadly speaking, sedimentary rocks accumulate in only a few types of plate tectonic settings, including passive continental margins, active continental margins, and volcanic island arcs.  The words passive and active refer to tectonic activity, implying regions of tectonic quiescence and region undergoing significant tectonic upheaval, respectively; volcanic island arcs are considered active by their very nature.  In the previous discussion of siliciclastic- and carbonate-dominated sedimentary systems, passive continental margins were correlated with the distal end of long siliciclastic systems, described by the Simple Ideal model, and with carbonate dominated systems; active continental margins and volcanic island arcs were correlated with the proximal end of some long siliciclastic systems, but more often with short siliciclastic systems (Figure 16 and Figure 25).  For a general description of the causes and processes associated with plate tectonics, and a more detailed description of tectonic settings, I refer my readers to the Plate Tectonics section of GEOLOGY BASICS.

Passive continental margins form at the trailing edges of continental masses that have long since broken apart and have rifted away from each other.  The continental margin gradually subsides over time as it accumulates a thick sequence of sedimentary rocks (Figure 26).  During initial rifting of the continent, a central axial graben forms at the incipient divergent boundary, consequent with adjacent horst mountains produced by upwarping of the crust and intermediate volcanism.  A great variety of sedimentary rocks are deposited within the graben associated with short system environments (Figure 16b) where lateral and vertical facies changes occur rapidly. Weathering and erosion of the highlands sheds coarse clastic feldspathic breccias and conglomerates into the graben to accumulate as alluvial fans, which grade basinward into braided stream and lake deposits; and these siliciclastics are often interbedded with intermediate composition lava flows and pyroclastic material.  Rift valley lakes are usually deep, alkaline, and anoxic, accumulating dark, organic-rich mudstones.  Complete rifting leads to outpouring of mafic flood basalts and eventual marine flooding of the elongate graben basin.  Conglomeratic fan deltas develop around the basin margin, while turbidity currents carry immature clastics into the deepening basin to form proximal submarine fans and more distally, dark laminated mudstones.  I refer to this diverse package of volcanics and siliciclastic sediments as a rift facies.  Over time, the significant relief of the rift zone becomes increasingly level as the horst mountains are removed by erosion and the axial graben, now a passive divergent continental margin, accumulates sediment and subsides (Figure 26).  Depositional control transitions from a short siliciclastic system to the distal portion of a long siliciclastic system, the transgressive beach sandstones, near shelf mudstones, and far shelf carbonates of the Simple Ideal model.  Eventually, continental terrain of any relief is far removed from the passive margin, siliciclastic input is much diminished or nonexistent, and a carbonate-dominated system takes over on the continental shelf and slope (Figure 26).  A carbonate shelf barrier reef and platform complex will persist and grow basinward as long as sea level remains stable or carbonate deposition keeps pace with subsidence and/or sea level rise.

Figure 26.  Typical sedimentary facies associations of a passive continental margin.

Active continental margins form where two tectonic plates undergo collision at a convergent plate boundary.  Volcanic island arcs form where two oceanic plates collide; the older, colder, denser oceanic plate subducts beneath the other and its partial melting provides the material that constructs the arcuate chain of volcanic islands.  Weathering and erosion of the islands dumps immature sediments corresponding to short system depositional environments into basins to either side of the volcanic arc.  Figure 27 provides a cross-sectional model of an oceanic-to-oceanic collision, its volcanic island arc, and associated arc sediment deposition.  As subduction initiates, early generated mafic magmas erupt on the sea floor to form pillow basalts and volcaniclastic rocks.  Broadening convergence and progressive heating of the subducting slab causes fractional melting of minerals lower on Bowen’s Reaction Series (Figure 9 in GEOLOGY BASICS) to produce magmas of intermediate composition.  The magmas rise and are emplaced into the growing arc, in time forming composite volcanoes on the sea floor that eventually build above sea level to create the island chain characteristic of the volcanic island arc.  The area on the subduction side of the arc forms the fore-arc; the area to the rear of the arc forms the back-arc.   The volcanic islands of the arc come under attack by weathering and erosion as soon as they breach the ocean surface and lithic sediments wash into fore-arc and back-arc basins.

Figure 27.  A cross-sectional model of a volcanic island arc formed by oceanic-to-oceanic collision and its associated arc-sediment deposition. 

In the back-arc, turbidity currents transport sediment onto the ocean floor to form submarine fans comprised of repeated Bouma sequences known as turbidites that over time, prograde basinward and coarsen upsection.  Sedimentary facies closer to the islands correspond to lithic conglomerate alluvial-fan environments that grade directly into lithic conglomerate gravel beaches, and then into steep, proximal submarine-fan environments.  Warm climates can produce fringing carbonate reefs.  On the reef’s seaward side, wave action produces bioclastic debris that mixes with lithic sediments cascading downslope; on the island side, oomicrite and biomicrite muds, and biosparite sands are deposited in a lagoonal environment.  Volcanic eruptions produce lava flows and pyroclastic material that may become interbedded with these deposits.  The fore-arc area contains all the features of the back-arc, with additional complexities of its own.  During early stages of arc development, the fore-arc basin surface descends rapidly, directly into the oceanic trench formed over the site of initial subduction.  The trench fills with turbidite deposits which are gradually dragged downward by subduction, to be folded, sheared, and combined with slivers of oceanic crust scraped up from the descending plate into mélange, and thrust upward onto the leading edge of the arc as an accretionary wedge.  As the wedge grows, a trough forms between it and the volcanic arc, formally called the fore-arc basin, which accumulates submarine fan turbidites during later stages of arc development.  The accretionary wedge may eventually grow to significant size, such that it to rises above the waves to form part of the island arc, only to contribute sediments to the fore-arc basin and oceanic trench to either side.    

Where an oceanic and continental plate collide, the denser oceanic plate subducts beneath the continental plate to form an active continental margin; partial melting of the subducting plate provides the material to construct a volcanic arc on the leading edge of the continental plate.  The tectonic setting is similar to the island arc generated by oceanic-to-oceanic collision shown in Figure 27, producing nearly all of the same geologic features, except the chain of volcanoes and back-arc area are perched on dry land (in mature active continental margins, the fore-arc area and accretionary wedge have risen above sea level as well).  In some instances, an oceanic plate carrying an island arc, or another continental plate altogether, can collide with a passive continental margin to produce a foreland basin setting.  Figure 28 depicts a continent-continent collision (although the collision could involve an island arc and a continental margin); in this situation neither continental plate will subduct because of their low density, but the active margin of one plate will partially slide up onto the passive margin of the other plate.  The overriding plate is referred to as the hinterland, the overridden plate is the foreland, and the region of collision is the suture zone.  During collision, the former active continental margin undergoes dramatic change.  The accretionary wedge of the active continental margin is squeezed from behind and from the fore, sheared, shortened, and thrust over the passive margin.  The volcanic arc migrates into the hinterland, but volcanic activity eventually ceases when subduction stops.  Volcanic rocks and crystalline basement of the continental plate are thrust toward the foreland as large, hinterland-dipping slabs as the mountain belt grows in height and width.  In the foreland region, the thick wedge of passive margin sedimentary rocks is compressed, folded, and thrust faulted toward the foreland.  The overriding active continental margin adds substantial mass to the passive continental margin, causing subsidence of the foreland.  A deep-water basin rapidly develops where the passive margin once occurred, and erosion of the hinterland mountains transports huge quantities of sediment basinward.  A thick clastic wedge progrades into the foreland basin comprised of a coarsening upward sequence, initially dominated by short system environments while the hinterland mountain belt is rising, but evolving over time into long system environments as the hinterland mountain belt is worn down by weathering and erosion.  Short system deposits consist of deep basin, dark, organic-rich mudstones, submarine fan turbidites (wacke sandstones and mudstones), gravel beach lithic to feldspathic conglomerates, and alluvial fan lithic to feldspathic breccias and conglomerates; long system sediments are comprised of wave- or tide-dominated shelf mudstones, beach and nearshore quartz sandstones, and meandering river feldspathic wacke sandstones.

Figure 28.  Continent-to-continent collision and the formation of a foreland basin.