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Geologic Maps:
A geologic map portrays the distribution of rocks, deposits, or other geologic features in a
specified area. Each consolidated rock type that can be distinguished by similar characteristics is
categorized into a mappable unit, or formation. Unconsolidated deposits such as landslides and
stream alluvium also are designated on our geologic maps. Unique colors, patterns, and labels
are used to differentiate each unit on the map. Colors are chosen by the age of the rocks being
described. For example, rocks from the Jurassic Period are colored in shades of green, and
Quaternary deposits are colored shades of yellow. Labels designate the age and name of the
formation or deposit. A geologic map is typically printed over a topographic base map. Many
different types of lines and symbols are found on geologic maps. The most prevalent are thin
black lines that depict the contacts between two different mappable units. Line width and color
are used to differentiate other types of features such as faults and folds. Geologists collect
structural information describing whether the rock layers are tilted or not. Most sedimentary and
volcanic rocks were originally deposited horizontally. Therefore, the tilt (dip) of layered rock
units may provide key information to understanding whether non horizontal units were deformed
by faulting or folding. This type of structural information also helps geologists build interpretive
cross-sections that depict how the map units may look in the subsurface. Geologic maps provide
a wealth of information to all who use them. They are important for geologic hazard detection
and mitigation, mineral and groundwater resource evaluation, and provide enjoyment for the
casual roadside geologist. Geologic maps help us understand the earth on which we walk, and
give us a greater appreciation for the geology around us.
Geological data is 3-dimensional; as exploration geophysicists you are interested in the
arrangement of rocks below the ground. Although seismic, magnetic and gravity data, among
other techniques, will tell you much of what you need to know, all clues help.
Before we go any further, we need to define some simple geological terms: Think about
sedimentary rocks (those that form in orderly layers). If the layers of rock are planar
(horizontal), with constant thickness and continue forever, then these rocks are said to have
layer-cake stratigraphy (This is rare in practice, but useful for now to help visualise these
ideas). We deal with structures in terms of the orientation of the planes (the boundaries between
units or beds. Bed is an informal term referring to individual, often relatively thin layers within
the rock. Loosely, each bed formed over a short period of time; often, the surface of a bed formed
the sediment surface at some point in the past. Units are a collection of adjoining beds that are
grouped together when they have some similarity e.g. mineralogy, palaeontology or particular
structures that indicate a common process in their origin. Units must be mappable and distinct
from one another, but the contact does not have to be particularly distinct e.g. “when the
sandstone component exceeds 75%”. Units are grouped together in stratigraphy as formations
and members of formations.
What do the outcrop patterns of these beds tell us?
Outcrop patterns represent the intersection of the 3-dimensional shape of the rock with the land
surface. Where the rocks are flat and the land is not, then boundaries will outcrop along
topographic contour lines. Vertical features and nothing else will outcrop as straight lines in hilly
terrain. All other orientations will wave across the landscape, cutting across topographic
contours. Understanding the scale and coordinates of a geological map Maps in general come in
a variety of scales depending on the amount of detail needed for different purposes. In all cases,
they will show a grid for location (in Britain, the National Grid on land), almost always with
North at the top. Points on the map are referred to using coordinates (eastings then northings)
which are usually 6 or 8 figure references. The basic grid square covers 100,000 metres, with
northings. Four figure grid references specify locations to within 10 metres. Maps covering a
larger area tend to use latitude and longitude in addition or instead. Maps for other countries will
use other National Grids.
What is strike and dip?
Geologists define the orientation of dipping beds using the terms strike and dip. Strike is the
azimuth (bearing on a compass) of a horizontal line on a bed i.e. a line perpendicular to the
steepest angle of dip). This is always given as a 3 figure number, e.g. 090 for a bed striking
East-West. Dip is the angle from the horizontal of the steepest gradient of the bedding surface.
This is a 2 figure number (e.g. a horizontal bed has a dip of 00 degrees ). It has an orientation and
dip arrows always point down dip. The two combined are given as 256/45 SE, for example, fully
defining the orientation of a plane. Strike and dip are generally marked on maps, using a
combination of symbols and figures, these are given below. Strike and dip are measured with a
compass/clinometer, on an area about 10cm x 10 cm. For convenience, we approximate even
wavy beds as planes. You can assess in the field how planar the beds are and ensure that you take
enough readings.
More on outcrop patterns
Outcrop patterns also provide information on the dip direction of simple sedimentary units.
Imagine the stratigraphy of an area as a pile of books. Let’s start with the books horizontal (they
stay on the desk better that way!), so we are considering rocks in which the bedding plane is
horizontal. How these beds look on a map depends on the landscape of the area. If we were in a
flat featureless area, then all we would see at the surface would be on type of rock; This would
be whichever rock happened to be at the same height as the plain. Therefore, the geological map
would show one type of rock only. But if the area had hills and valleys that cut through the
horizontal layers of rock, then we would see more than one rock type (i.e. more than one colour
of rock on the map). If the map has topographic contours on it, then the boundaries between rock
types will outcrop in shapes that follow the contours (because the beds are horizontal, the
boundaries between them can exist only at particular heights. That is the height at which they
outcrop. So, horizontal beds produce distinctive outcrop patterns. The pattern is also clear if the
map is small scale; This is because flat beds produce elaborate, crenulated outcrop patterns that
reflect small details of the landscape. Look out for outcrops that zig-zag across streams, for
example. The other extreme comes from areas with vertical beds. Here the books are lined up as
if on a shelf. Vertical beds can outcrop only as straight lines, completely unaffected by the
topography. The general case is when the rocks are dipping. They outcrop along lines that do not
follow the contours, nor are they straight lines. They mark the intersection of the shape of the
boundaries between beds, with the landscape surface in three dimensions. Gently dipping beds
outcrop in patterns that are almost the same as topographic contours, but the bed boundaries
cross them occasionally. Steeply dipping beds outcrop in lines that are close to straight, cutting
across many topographic contours. The shape of the outcrop can help you assess the dip of the
bed and especially the relative dips of different beds in the same landscape.
Dip direction from outcrop
We can understand the direction of dip if we know the relative ages of the rocks. Look at the
stratigraphic column at the side of a map. This shows the units visible on the map in their order
of formation (oldest at the bottom, youngest at the top). Returning to the pile of books concept;
where the units are tilted, the order in which they appear on the map tells you the direction of
dip: Dip direction is from older to younger. (Please note: This won’t work in areas of complex
structure, where beds may be upside down!) Outcrop patterns will change in response to changes
in dip, thickness and of course, the topography of the land surface.
True and apparent dip
True dip is a line perpendicular to the strike and is the steepest line along the plane of the bed.
Whereas apparent dip is the angle from a horizontal line that is not perpendicular to the strike. (If
you find this difficult to visualise there is a 3D model at the back of this document you can
Dip variations
Folds and faults are the most common causes of variation in strike and dip; see attached sheets
for brief examples. Folding and faulting, followed by subsequent erosion and deposition of a
younger rock produces an unconformity (see figure below). Unconformities are variations to the
simplest case of sedimentary rocks. They are identifiable on maps as place where more than one
younger rock is in contact with several older rocks (in the simplest case). Unconformities make
useful time markers.
But in many cases, sedimentary rocks are not of uniform thickness, nor do they continue the
same forever. In some cases, the stratigraphic column on the side of a map specifies bed
thicknesses and variations in the type of rock of a particular age (facies variations) are also
shown in some cases. In cases where sediment thicknesses are known to vary significantly, a
special type of map is used to illustrate this (an isopach or isopachyte map), which show
contours of bed thickness. These are very useful in sedimentary basin analysis, where sediment
thickness is an important variable for modelling.
Calculating bed thickness
True thickness (t) is measured perpendicular to upper and lower surface of the bed. V is vertical
thickness that would be encountered in a borehole. The true thickness of a bed can be obtained
from the outcrop width (w) and dip as follows:
True thickness (t) = width of outcrop (w) x sinθ (angle of dip)
Synthesis of geological maps
The data presented on even simple geological maps can be very confusing. We need to simplify
it, in order to answer whatever questions we have about the area. There are two ways to do this,
which may be appropriate in different cases:
● Cross section (precise or sketch)
● Sketch map
Begin by dividing the map up into geologically sensible regimes (use faults and unconformities
to define domains on the map). Put these boundaries onto your sketch map (as well as grid refs
and scale). Then, look at the map and deduce the orientation of the rocks in each area: Which
way do they dip? Are they steep or shallow? Are there folds? Are there important faults? Etc.
Use a key and colour or shade your map to show important groups of rocks. Show the dip
direction and amount for each domain.
Hints and tips for sketch/summary maps
A good sketch map:
● Is the same shape as the original map
● Is easy to relate to the original map, i.e. it uses similar colours and ornament, if needed,
and has grid lines, key and scale shown clearly.
● Is a sketch, i.e. it reproduces the general form of key boundaries without following every
twist and turn on the map.
● Reproduces accurately, but not in full detail, significant outcrop shapes, boundaries and
cross-cutting relationships.
● Gives indications of topography, for example showing spot heights on high ground and
major rivers to show valleys.
● Summarises complex stratigraphy, for example by grouping units with similar histories
together and treating them as one unit.
● Highlights important features such as major faults and folds by marking them and
showing as much information deduced from the map overall as possible. For example,
folds can be shown as axial traces, synformal or antiformal, with younging directions
marked. Faults traces should show dip direction, downthrow side and throw (if known),
marked according to convention.
● Is annotated, to indicate complexities not easily summarised, e.g. “region of many
parallel steep faults” or “many thin bedded-parallel igneous intrusions”.
● Shows representative dip directions and dip amounts where feasible.
● Is neat and easy to read.
A bad sketch map:
● Has a shape and form different from the original map.
● Does not have grid lines numbered, has no scale, and it is hard to tell which bit of it refers
to what part of the original map.
● Uses colours and ornament that do not relate to those of the original map, uses too many
colours or uses ornaments without a key.
● Does not have a key.
● Is a detailed reproduction of the original map, right down to the drift in the stream
valleys. (If you want a smaller reproduction of the map, use a photocopier!)
● Retains all the stratigraphic detail of the original.
● Does not highlight the important features of the map as a whole: any analysis of this
small map is as difficult as, if not harder than, analysis of the original.
● The idea is to sort out what is important and show these features on your sketch.
● Does not give any clue to topography.
● Does not show dip or even dip directions.
● Has no helpful annotation.
● Is messy, has imprecise line work and is generally hard to read.
Cross-sections show the thicknesses, dip directions and relationships between units on a vertical
slice through a map. They are a useful way of synthesising data. Accurate sections match exactly
with the rocks that outcrop along the section line. Sketch sections sacrifice precision for a clear
representation of the structure and often combine information from across the map onto a section
to make it representative of the area as a whole. Cross-sections generally are most useful to
geologists when the horizontal scale equal to the vertical one (this means that dip amounts on the
section are accurate). Any form of vertical exaggeration alters dips and can produce misleading
structures. Both types of cross-section will show clearly such features as unconformities, folds,
faults, sediment thickness changes, igneous intrusions.
A sketch map is a simplified diagram of the map, highlighting important information and
neglecting both detail and superfluous aspects of the geology. They will look very different
depending on their purpose. In general, such sketch maps will also highlight features such as
unconformities, faults, folds and igneous intrusions. But they may also illustrate features such as
fault density, mineralisation or porosity.
Structure contours
Structure contours are simply contours of height drawn on a particular geological surface (e.g. a
bedding plane, a coal seam or a fault). They are usually drawn as height above some datum or
reference height, such as sea level. They are a clear way of representing what may be a complex
shape in the rock underground. You draw them on the basis of the information that is available
(often depths to a particular layer from boreholes or seismic sources).
In summary:
● If the structure contours are straight, parallel and equidistant, then the surface is planar.
Its strike is constant and parallel to the strike of the structure contours.
● If structure contours are curved, then the strike of the bed varies.
● If the separation of the structure contours varies, the dip of the bed varies.
When rocks are folded, they also assume typical outcrop patterns, as shown in figure 2 below
(characteristically forming V-shapes in valleys and ridges).
Faults are surfaces in the Earth across which there has been some displacement, usually by
cataclasis (the deformation of rock via crushing and shearing). Faults are usually narrow in
proportion to their length and breadth, often planar or gently curved and exist mainly in the top
10-15 kilometres of the Earth’s crust. Below this depth, rock deforms in a plastic fashion,
without fracturing. Because faults involve displacement, one of the targets of geologists is to
quantify this displacement (ideally as a vector). To do this, they need a unique marker that can be
identified in the rocks on each side of the fault (this is rare). It is usual for faults to offset beds
(i.e. planar features). This is not enough information to determine the movement of the fault. It
can tell you the offset on the fault (the separation between units that were once continuous).
Offset is described in terms of the horizontal offset, called heave; and the vertical offset called
throw (see figure 3). There are many circumstances in which you cannot tell heave or throw (e.g.
if horizontal beds are displaced horizontally, heave and throw are both zero. Consider the
example of dipping beds with measurable heave and throw; displacement could be solely
horizontal, solely vertical or oblique.
Faults- places where Earth’s crust moved, Shown by thick black line, Three types of faults:
Normal, Reverse, Strike-slip
● Two types of Strike- slip faults: dextral and sinistral
Dip-slip faults are inclined fractures
where the blocks have mostly shifted
vertically. If the rock mass above an
inclined fault moves down, the fault is
termed normal, whereas if the rock
above the fault moves up, the fault is
termed reverse. A thrust
fault is a reverse fault with
a dip of 45 degrees or less.
Units- different types of rocks/minerals
● Description of Map Units
○ Shows color, name abbreviation and physical
characteristics of map unit
● Correlation of Map units
○ Show relative age of each material
■ Contacts- show where one type of rock passes into another
● Clear transition- solid line
● Unclear/ Approximate transition- dashed line
Strike and Dip- Strike l plane. Dip is the angle at which a planar feature is inclined to the
horizontal plane; it is measured in a vertical plane perpendicular to the strike of the feature. A
bed will be used in this description. For a recently created flat bed it would have a dip of 0° and a
directionless strike. When it is tilted it gains a direction. The bed is usually not homogeneous
since it usually folds when tilting, fractured, or compressed. A strike of 000° means the bed is
dipping east; 090° for south; 180° for west; and 270° for north. A dip of 0° means its flat and 90°
for a vertical bed. Trend is the direction of the line formed by the intersection of the planar
feature with the ground surface; trend is the same as strike only if the ground surface is parallel
to the horizontal plane. Plunge is the vertical angle between the horizontal plane and the axis or
line of maximum elongation of a feature.
The first thing that needs to be done is to calibrate your compass to the magnetic declination.
Magnetic declination is the difference between geographic north and magnetic north, with
respect to your position. To determine what degree you need to
set your compass to consult the picture below. To find the strike,
stand so that the dip of the bed slopes to the right (this is known
as the Right Hand Rule). Orient the long arm of the Brunton in
the direction you are facing, and place the side of the Brunton
against the rock. Level it so that the bubble in the flat level is in
the center of its circle, then read the number that the north end
of the needle is pointing at. That number is your strike. To find
the dip, orient your Brunton perpendicular to the strike. The
easiest way to do this is to let water fall on the surface of the
rock and it tends to roll down the steepest angle. Place the
Brunton against the rock, on the side this time, with the long
arm pointed down the dip of the bed. It may be complicated
sometimes if a rock surface is rough. To deal with this you put a
flat object, like a clipboard, on the rock to determine its strike
and dip. Do note that if the surface is weathered too much then it
may not be the true dip of that part of the bed; so you would
need to search other places along the bed to find a reasonable
outcrop. Record the data of a bed with a strike of 15 and dip of
45 as 015/45. This can be very useful when figuring out if the area is influenced by a thrust fault,
normal fault, or strike slip fault through the use of a focal mechanism diagram.
We can express the orientation of a bed (or any other planar feature) with two values: first, the
compass orientation of a horizontal line on the surface—the strike—and second, the angle at
which the surface dips from the horizontal, (perpendicular to the strike)—the dip (Figure 12.18).
It may help to imagine a vertical surface, such as a wall in your house. The strike is the compass
orientation of the wall and the dip is 90˚ from horizontal. If you could push the wall so it’s
leaning over, but still attached to the floor, the strike direction would be the same, but the dip
angle would be less than 90˚. If you pushed the wall over completely so it was lying on the floor,
it would no longer have a strike direction and its dip would be 0˚. When describing the dip it is
important to include the direction. In other words. if the strike is 0˚ (i.e., north) and the dip is 30˚,
it would be necessary to say “to the west” or “to the east.” Similarly if the strike is 45˚ (i.e.,
northeast) and the dip is 60˚, it would be necessary to say “to the northwest” or “to the
Measurement of geological features is done with a special compass that has a built-in clinometer,
which is a device for measuring vertical angles. An example of how this is done is shown on
Figure 12.19.
● Shows orientation of rocks and fault planes
● Strike- long line
○ direction the rocks point
● Dip- short line with number
○ Direction and angle rocks are tilted
■ Cross Sections- vertical slice showing layers of rock
Plate Tectonics
From the deepest ocean trench to the tallest mountain, plate tectonics explains the features and
movement of Earth's surface in the present and the past.
Plate tectonics is the theory that Earth's outer shell is divided into several plates that glide over
the mantle, the rocky inner layer above the core. The plates act like a hard and rigid shell
compared to Earth's mantle. This strong outer layer is called the lithosphere, which is 100 km (60
miles) thick, according to Encyclopedia Britannica. The lithosphere includes the crust and outer
part of the mantle. Below the lithosphere is the asthenosphere, which is malleable or partially
malleable, allowing the lithosphere to move around. How it moves around is an evolving idea.
Developed from the 1950s through the 1970s, plate tectonics is the modern version of
continental drift, a theory first proposed by scientist Alfred Wegener in 1912. Wegener didn't
have an explanation for how continents could move around the planet, but researchers do now.
Plate tectonics is the unifying theory of geology, said Nicholas van der Elst, a seismologist at
Columbia University's Lamont-Doherty Earth Observatory in Palisades, New York.
"Before plate tectonics, people had to come up with explanations of the geologic features in their
region that were unique to that particular region," Van der Elst said. "Plate tectonics unified all
these descriptions and said that you should be able to describe all geologic features as though
driven by the relative motion of these tectonic plates."
There are nine major plates, according to World Atlas. These plates are named after the
landforms found on them. The nine major plates are North American, Pacific, Eurasian, African,
Indo-Australian, Australian, Indian, South American and Antarctic.
The largest plate is the Pacific Plate at 39,768,522 square miles (103,000,000 square kilometers).
Most of it is located under the ocean. It is moving northwest at a speed of around 2.75 inches (7
cm) per year.
There are also many smaller plates throughout the world.
e driving force behind plate tectonics is convection in the mantle. Hot material near the Earth's
core rises, and colder mantle rock sinks. "It's kind of like a pot boiling on a stove," Van der Elst
said. The convection drive plates tectonics through a combination of pushing and spreading apart
at mid-ocean ridges and pulling and sinking downward at subduction zones, researchers think.
Scientists continue to study and debate the mechanisms that move the plates.
Mid-ocean ridges are gaps between tectonic plates that mantle the Earth like seams on a baseball.
Hot magma wells up at the ridges, forming new ocean crust and shoving the plates apart. At
subduction zones, two tectonic plates meet and one slides beneath the other back into the mantle,
the layer underneath the crust. The cold, sinking plate pulls the crust behind it downward.
Many spectacular volcanoes are found along subduction zones, such as the "Ring of Fire" that
surrounds the Pacific Ocean.
Subduction zones, or convergent margins, are one of the three types of plate boundaries. The
others are divergent and transform margins.
At a divergent margin, two plates are spreading apart, as at seafloor-spreading ridges or
continental rift zones such as the East Africa Rift.
Transform margins mark slip-sliding plates, such as California's San Andreas Fault, where the
North America and Pacific plates grind past each other with a mostly horizontal motion.
While the Earth is 4.54 billion years old, because oceanic crust is constantly recycled at
subduction zones, the oldest seafloor is only about 200 million years old. The oldest ocean rocks
are found in the northwestern Pacific Ocean and the eastern Mediterranean Sea. Fragments of
continental crust are much older, with large chunks at least 3.8 billion years found in Greenland.
With clues left behind in rocks and fossils, geoscientists can reconstruct the past history of
Earth's continents. Most researchers think modern plate tectonics began about 3 billion years
ago, based on ancient magmas and minerals preserved in rocks from that period. Some believe it
could have started a billion years after Earth's birth, at around 3.5 billion years.
"We don't really know when plate tectonics as it looks today got started, but we do know that we
have continental crust that was likely scraped off a down-going slab [a tectonic plate in a
subduction zone] that is 3.8 billion years old," Van der Elst said. "We could guess that means
plate tectonics was operating, but it might have looked very different from today."
As the continents jostle around the Earth, they occasionally come together to form giant
supercontinents, a single landmass. One of the earliest big supercontinents, called Rodinia,
assembled about 1 billion years ago. Its breakup is linked to a global glaciation called Snowball
A more recent supercontinent called Pangaea formed about 300 million years ago. Africa, South
America, North America and Europe nestled closely together, leaving a characteristic pattern of
fossils and rocks for geologists to decipher once Pangaea broke apart. The puzzle pieces left
behind by Pangaea, from fossils to the matching shorelines along the Atlantic Ocean, provided
the first hints that the Earth's continents move.
Plates bumping into each other can also cause mountain ranges. For example, India and Asia
came together about 55 million years ago, which created the Himalaya Mountains, according to
National Geographic.
Plate Boundaries:
1. Convergent boundaries: where two plates are colliding.
Subduction zones occur when one or both of the tectonic plates are composed of oceanic crust.
The denser plate is subducted underneath the less dense plate. The plate being forced under is
eventually melted and destroyed.
i. Where oceanic crust meets ocean crust
Island arcs and oceanic trenches occur when both of the plates are made of oceanic crust. Zones
of active seafloor spreading can also occur behind the island arc, known as back-arc basins.
These are often associated with submarine volcanoes.
ii. Where oceanic crust meets continental crust
The denser oceanic plate is subducted, often forming a mountain range on the continent. The
Andes is an example of this type of collision.
iii. Where continental crust meets continental crust
Both continental crusts are too light to subduct so a continent-continent collision occurs, creating
especially large mountain ranges. The most spectacular example of this is the Himalayas.
2. Divergent boundaries – where two plates are moving apart.
The space created can also fill with new crustal material sourced from molten magma that forms
below. Divergent boundaries can form within continents but will eventually open up and become
ocean basins.
i. On land
Divergent boundaries within continents initially produce rifts, which produce rift valleys.
ii. Under the sea
The most active divergent plate boundaries are between oceanic plates and are often called
mid-oceanic ridges.
3. Transform boundaries – where plates slide passed each other.
The relative motion of the plates is horizontal. They can occur underwater or on land, and crust
is neither destroyed nor created.
Because of friction, the plates cannot simply glide past each other. Rather, stress builds up in
both plates and when it exceeds the threshold of the rocks, the energy is released – causing
Rock Formation
Igneous rocks are one type of rock. These rocks are associated with volcanoes and form at plate
boundaries, either as magma under the ground hardens or as lava flows over the surface and
cools. Magma cools to form intrusive igneous while lava cools to form extrusive igneous rocks.
Many igneous rocks are basalt or granite, two of the most abundant rock types on the planet.
Other examples of igneous rocks include andesite, rhyolite, granodiorite and gabbro.
Sedimentary rocks take thousands of years to form. Unlike igneous rocks which can form as the
result of violent collisions and volcanic eruptions, sedimentary rocks form quietly, as sand, mud
and sometimes the remains of living things collect on the sea floor or on land. As more and more
sediment deposits, the sheer weight of the sediments causes them to compress, forming solid
rock. Sedimentary rocks are distinguished by their layered appearance, as different types of
sediments collect over time, and by the presence of fossils. Examples of sedimentary rocks
include conglomerate, sandstone, mudstone and limestone.
Metamorphic rocks are the result of pressure and heat applied to igneous or sedimentary rocks.
The appearance of these rocks is transformed along with the structure; for example, metamorphic
sedimentary rocks retain the layers, but the layers are bent and compressed. Some examples of
metamorphic rock include marble, slate and gneiss.
The most beautiful jewelry often features gemstones, and most gemstones come from
underground rock formations. Pressure and heat cause fluids to crystallize in different ways,
leading to the formation of such gemstones as onyx, ruby, sapphire and turquoise.
The Earth is a closed system, which means that what is formed on the planet must be recycled
and re-used. Just as the water cycle explains how water is cycled through the atmosphere, the
rock cycle shows how rocks are created and destroyed. The cycle usually begins with the
emplacement or eruption and eventual erosion of igneous rocks and their subsequent deposition
as sediments. From here the rocks may be pushed further into the Earth and incorporated into the
molten inner layers, or they may be brought to the surface over time as sedimentary rocks. The
rock cycle is multi-directional, so the formation of rocks is continuous but happens in different
ways over time.
In most areas of the Earth, rocks are being constantly created or destroyed. However, igneous,
sedimentary and metamorphic rocks may be millions of years old. The oldest rock formation is
believed to be in Western Australia, a combination of metamorphic and sedimentary rocks
known as Jack Hills. They have been dated at 4.4 billion years old, forming only a few million
years after the Earth itself.
Synclines and Anticlines
Syncline and anticline are terms used to describe folds based on the relative ages of folded rock
layers. A syncline is a fold in which the youngest rocks occur in the core of a fold (i.e., closest to
the fold axis), whereas the oldest rocks occur in the core of an anticline. Synclines form when
rocks are compressed by plate-tectonic forces. They can be as small as the side of a cliff or as
large as an entire valley.A fold of rock
layers that slope downward on both
sides of a common crest. Anticlines
form when rocks are compressed by
plate-tectonic forces.
Basins- A basin is a depression, or
dip, in the Earth’s surface. Basins are
shaped like bowls, with sides higher
than the bottom. They can be oval or
circular in shape, similar to a sink or
tub you might have in your own
bathroom. Some are filled with water.
Others are empty. Basins are formed by forces above the ground (like erosion) or below the
ground (like earthquakes). They can be created over thousands of years or almost overnight. The
major types of basins are river drainage basins, structural basins, and ocean basins.
○ River Drainage Basins
■ A river drainage basin is an area drained by a river and all of its tributaries
■ Made up of many different watersheds.
■ Watershed is a small version of a river basin. Every stream and tributary
has its own watershed, which drains to a larger stream or wetland. These
streams, ponds, wetlands, and lakes are part of a river basin.
■ Every river is part of a network of watersheds that make up a river
system’s entire drainage basin. All the water in the drainage basin flows
downhill toward bigger rivers.
○ Structural Basins
■ Structural basins are formed by tectonic activity (movement of large
pieces of the Earth’s crust, called tectonic plates)
■ The natural processes of weathering and erosion also contribute to forming
structural basins.
● Structural basins form as tectonic plates shift. Rocks and other
material on the floor of the basin are forced downward, while
material on the sides of the basin are pushed up. This process
happens over thousands of years.
● If a basin is shaped like a bowl, a structural basin is shaped like a
series of smaller bowls, stacked inside each other.
● Usually found in dry regions.
● Some structural basins = endorheic basins.
○ Have internal drainage systems = don’t have enough water
to drain to a stream, lake, or ocean
water that trickles into these types of basins
evaporates or seeps into the ground.
○ When enough water collects in an endorheic basin,forms a
very salty lake. While water evaporates into the
atmosphere, minerals remain. The remaining water
becomes even saltier.
○ A lake basin is another type of structural basin
■ Form in valleys blocked by rocks or other debris
left by a landslide, lava flow, or glacier.
■ Debris acts = dam, trapping water and forming a
■ Lake basins may also be carved out by glaciers as
they move down valleys or across the land. When
the glaciers move, the basins they create remain.
■ Sedimentary basins are a type of structural basin
that aren’t shaped like typical basins, sometimes
forming long troughs.
■ Sedimentary basins have been filled with layers of
rock and organic material over millions of years.
■ Material that fills up the basin is called sediment
■ Key sources of petroleum and other fossil fuels.
○ Ocean Basins
■ Ocean basins are the largest depressions on Earth. Edges of the continents,
called continental shelves, form the sides of ocean basins.
■ 5 major basins: he Pacific basin, the Atlantic basin, the Indian basin, the
Arctic basin, and the Southern basin
■ Smaller basins are oceanics basins such as the North Aleutian Basin,
between the Pacific and Arctic Oceans.
■ Tectonic activity constantly changes ocean basins. Seafloor spreading and
subduction are the most important types of tectonic activity that shape
ocean basins.
■ Seafloor spreading: happens along the boundaries of tectonic plates that
are moving apart from each other. These areas are called mid-ocean
ridges. New seafloor is created at the bottom, or rift, of a mid-ocean ridge.
Ocean basins that have mid-ocean ridges are expanding. The Atlantic
basin, for instance, is expanding because of seafloor spreading.
■ Subduction happens along the boundaries of tectonic plates that are
crashing into each other. In these subduction zones, the heavier plate
moves underneath, or subducts, the lighter one. Ocean basins that
experience subduction, such as the Pacific basin, are shrinking.
● Types of Sedimentary Basins
○ Rift Basin: The down-dropped basin
formed during rifting because of stretching
and thinning of the continental crust
○ Passive Margins: Subsidence along a passive
margin, mostly due to long-term accumulation
of sediments on the continental shelf
○ Trench: Downward flexure of the subducting
and non-subducting plates (sites of
accretionary wedges)
○ Forearc Basin: The area between the
accretionary wedge and the magmatic arc,
largely caused by the negative buoyancy of the
subducting plate pulling down on the
overlying continental crust
○ Foreland Basin: A depression caused by the
weight of a large mountain range pushing the
adjacent crust below sea level
○ Strike-slip Basin: A
pull-apart block (eg. between two
transform faults) that subsides
significantly (ex. San Andreas Fault)
● Monoclines: Monoclines are
folds consisting of two horizontal (or
nearly so) limbs connected by a
shorter inclined limb. They can be
compared to anticlines, which
consist of two inclined limbs dipping
away from each other, and synclines,
which consist of two inclined limbs dipping towards each other.Folds
such as monoclines, anticlines, and synclines are defined solely on the basis of their
geometry, and the names therefore have
no genetic connotations. Monoclines are,
however, characteristic of regions in which sedimentary rocks have been deformed by dip
slip movement along vertical or steeply dipping faults in older and deeper rocks, such as
the Colorado Plateau of the southwestern United States.
○ 4 ways of formation
■ By differential compaction over an underlying structure, particularly a
large fault at the edge of a basin due to the greater compactibility of the
basin fill, the amplitude of the fold will die out gradually upwards.
■ By mild reactivation of an earlier extensional fault during a phase of
inversion causing folding in the overlying sequence.
■ As a form of fault propagation fold during upward propagation of an
extensional fault in basement into an overlying cover sequence.
■ As a form of fault propagation fold during upward propagation of a
reverse fault in basement into an overlying cover sequence.
Unconformities: the contact between sedimentary rocks that are significantly different in age or
between sedimentary rocks and older, eroded igneous or metamorphic rocks. Unconformities
represent gaps in the geologic record, periods of time that are not represented by any rocks.
There are three main types
○ Disconformities- beds of the rock sequence above and below the unconformity
are parallel to one another, but there is a measurable age difference between the
two sequences. The disconformity surface represents a period of non deposition
and/or erosion .
○ Angular Unconformities- At an angular unconformity, strata below the
unconformity have a different attitude than strata above the unconformity. Beds
below the unconformity are truncated at the unconformity, while beds above the
unconformity roughly parallel the unconformity surface.
○ Non Conformities- Nonconformity is used for unconformities at which strata
were deposited on a basement of older crystalline rocks. The crystalline rocks
may be either plutonic or metamorphic. For example, the unconformity between
Cambrian strata and Precambrian basement in the Grand Canyon is a
Geologic Principles:
The Law of Superposition states that beds of rock on top are usually younger than those
deposited below. This is logical, consider a layered cake or a stack of books, you can’t add
another layer unless one already exists to begin with. By understanding the Law of Superposition
we can make general statements about the ages of these rock units.Consider these top layers –
Unit K (dark green) is younger than Unit J (burnt orange) because it lies atop it, this also directly
relates to the relative age dating. In 1669 Nicolaus Steno made the first clear statement that strata
(layered rocks) show sequential changes, that is, that rocks have histories. From his work in the
mountains of western Italy, Steno realized that the principle of superposition in stratified
(layered) rocks was the key to linking time to rocks. In short, each layer of sedimentary rock
(also called a “bed”) is older than the one above it and younger than the one below it. Steno’s
seemingly simple rule of superposition has come to be the most basic principle of relative
dating. Steno originally developed his reasoning from observations of sedimentary rocks, but the
principle also applies to other surface-deposited materials such as lava flows and beds of ash
from volcanic eruptions.
In addition, Steno realized the importance of another principle, original horizontality, namely
that strata are always initially deposited in nearly horizontal positions. Thus, a rock layer that is
folded or inclined at a steep angle must have been moved into that position by crustal
disturbances (i.e., mountain building, faults, or plate tectonics) sometime after its deposition. The
Law of Original Horizontality suggests that all rock layers are originally laid down (deposited)
horizontally and can later be deformed. This allows us to infer that something must have
happened to the rocks to make them tilted. This includes mountain building events, earthquakes,
and faulting.The rock layers on the bottom have been deformed and are now tilted. The rock
layers on the top were deposited after the tilting event and are again laid down flat.
The Law of Lateral Continuity suggests that all rock layers are laterally continuous and may be
broken up or displaced by later events. This can happen when a river or stream erodes a portion
of the rock layers. This can also happen when faulting occurs. Faulting causes displacement in
rock units. The figure here shows the offset between the layers signified by the black line cutting
across the rocks. Trace the colors or letters across to find the layers that match. The rock layers
on the top seem to form a valley but we can tell that Unit I (dark blue) on one side is the same as
the Unit I (dark blue) on the other side. There is missing rock in between and a displacement
caused by deformation.
Cross-cutting relationships also helps us to understand the timing of events. Younger features
cut across older features. Going back to the fault on this image, we know that these rock layers
were involved in the fault movement because they are all offset. We can also determine which
beds of rock were tilted and that relationship to the rocks that are not tilted. The rock layers on
the bottom have been deformed and are now tilted. The rock layers on the top were deposited
after the tilting event and are again laid down flat. James Hutton’s observations related to
uniformitarianism also serve as the basis for another important geologic principle called
cross-cutting relationships, which is a technique used in relative age dating. In short an
intrusive rock body is younger than the rocks it intrudes. For example, Salisbury Crag, a
prominent Edinburgh landmark known to Hutton, owes its relief to a thick sheet of resistant
basalt. Hutton showed from the super-heated contacts below and above, and from places where
the basalt actually invaded underlying and overlying beds, that the thick basalt body was not
merely a flow that had formed in sequence. Rather it was intruded as hot magma into the
surrounding sedimentary rocks long after they were deposited (Eicher 1976). Other similar
relationships include faults being younger than the rock layers they cut and erosional surfaces
being younger than the rocks they erode.
The idea of Components is simple. If you find a rock that has other smaller pieces of rocks
within it, the smaller rocks inside must have existed before the larger rock was created.
The Principle of Faunal Succession states that a species appears, exists for a time, and then goes
extinct. Time periods are often recognized by the type of fossils you see in them. Each fossil has
a ‘first appearance datum’ and a ‘last appearance datum’. This is simply the oldest recorded
occurrence of a fossil and then the youngest recorded occurrence of a fossil. As an engineer and
surveyor, William Smith’s (1769–1839) travels acquainted him intimately with much of
England’s countryside. As he investigated roads, quarries, mines, and canals, Smith recognized
and traced out numerous sedimentary rock units, in which he noticed that each successive unit
contained its own diagnostic assemblage of fossils. Smith’s discovery that strata may be
identified by the fossils they contain became known as the law of faunal succession. Henceforth,
fossils became a new tool by which geologists could distinguish rock units of different ages from
one another. Faunal succession became a unifying principle by which rock units are categorized
and recognized widely.
This important principle raised questions about ancient life that were not easily answered at
Smith’s time, but even without answers to these questions, correlation between distant localities
now became feasible. A method was established to assign a stratigraphic classification based on
time relations of strata rather than on rock types, which was previously thought to indicate age.
In short, this was the key discovery that stratigraphic geology needed in order to progress further
(Eicher 1976).
Organic Evolution:
Organic evolution is the theory that more recent types of plants and animals have their origins in
other pre-existing forms and that the distinguishable differences between ancestors and
descendents are due to modifications in successive generations. Charles Darwin (1809–1882) did
not invent the idea of organic evolution; generations preceding him entertained the notion such
as the French zoologist Jean Baptist de Lamarck (1744–1829), a pioneer in invertebrate
paleontology, and Erasmus Darwin (1744–1802), grandfather of Charles. Until Charles Darwin’s
time, however, the idea had never had wide currency because earlier workers lacked important
data and the Huttonian concept of geologic time, which are both vital for the evolutionary
argument. Charles Darwin’s contribution was to propose a mechanism—natural selection—to
explain how this change could occur.
Because sufficient time is required to produce changes, one of the first requirements of evolution
came out of uniformitarianism. This idea provided the first link in Darwin’s chain of reasoning.
The relationship between organic evolution and geologic time is that the very randomness of
very small, fortuitous evolutionary variations in species, which result in noticeable changes in
the physical form of an organism, requires enormous amounts of time. Building on
uniformitarianism, Darwin constructed a singularly rational and convincing argument for the
origin of the diverse organisms that populate the world.
In his work, On the Origin of Species, first published in 1859, Darwin demonstrated the
existence of organic evolution to scientists and non-scientists alike. Geologic thought, indeed all
philosophical thought, has never been the same. Eicher (1976) provides the following summary:
● Populations of animals and plants produce progeny at such a rate that were they all to
survive, they would increase rapidly year after year.
● Spectacular progressive increases in population size do not, in fact, occur. Although most
populations fluctuate year by year, they remain essentially constant over the long term.
● A very real struggle for existence occurs in nature. Each individual must compete for
food and must cope successfully with every facet of the environment—both physical,
such as climate extremes, and biological, such as diseases and predators—in order to live
to produce progeny.
● Each individual differs from virtually all others in its species. By Darwin’s time, striking
variation in domestic animals had already been produced by selective breeding. Darwin
noted that species in nature had similar potential for modification.
● Here Darwin made a break with all previous suggestions on the subject. Instead of
postulating that modifications are induced by the environment and are then passed on
from generation to generation, he suggested that new characteristics arise from within an
organism entirely by chance. (We now know that these arise as genetic mutations.)
However, not all of these new characteristics will have adaptive significance or survival
value, and many may even be lethal.
● Some of the new characteristics enhance an individual’s success in coping with the
environment and may even allow the organism to push beyond previous environmental
barriers. Others will be unsuccessful, and individuals with these modifications will
simply not survive to pass them along; Darwin termed this process natural selection.
At the same time as Darwin, but unbeknown to him, another scientist, Gregor Mendel (1865),
demonstrated that organisms acquire traits through “discrete units of heredity,” later to become
known as genes. Much later, scientists in the 1930s showed that genes are the ultimate source of
variation within a population. That is, all variations arise through changes, called mutations, in
genes. If a mutation enables an organism to survive or reproduce more effectively, that mutation
tends to be preserved and spreads in a population through natural selection. Hence, evolution
depends on both natural selection and genetic mutations: mutations provide abundant genetic
variation, and natural selection sorts out the useful changes from the deleterious ones (Kennedy
et al. 1998).
With the discovery of DNA (deoxyribonucleic acid) in 1953 by Francis Crick and James Watson,
the study of evolution entered yet another phase, taking it to its most fundamental level. Crick
and Watson found that DNA contains the genetic instructions used in the development and
functioning of all known living organisms. Chemically DNA is a long polymer of simple units
called nucleotides, with a backbone made of sugars and phosphate atoms joined by ester bonds.
Attached to each sugar is one of four types of molecules called bases. It is the sequence of these
four bases along the backbone that encodes information, and the main role of DNA is the
long-term storage of information. Eukaryotic organisms such as animals, plants, and fungi store
their DNA inside the cell nucleus, while in prokaryotes such as bacteria it is found in the cell’s
cytoplasm. Within cells, DNA is organized into structures called chromosomes and the set of
chromosomes within a cell make up a genome. These chromosomes are duplicated before cells
divide, in a process called DNA replication.
DNA is the ultimate source of both change and continuity in evolution. The modification of
DNA through occasional changes or rearrangement in the base sequences underlies the
emergence of new traits, and thus of new species. At the same time, all organisms use the same
molecular codes of the four DNA base sequences. This uniformity in the genetic code is
powerful evidence for the interrelatedness of living things, suggesting that all organisms
presently alive share a common ancestor that can be traced back to the origins of life on Earth
(Kennedy et al. 1998).
Organic Extinction:
For more than 300 years, naturalists have recognized that fossils are the remains of once-living
plants and animals. They postulated that formerly existing species were no longer living; that is,
they had become extinct. However, the idea of extinction took a while to be proven to the
satisfaction of most people. The chief reason was that the best known fossils were marine
invertebrates. These are certainly as good as terrestrial organisms to demonstrate extinction, but
at the beginning of the 19th century little was known of life in the oceans. Mindful of the very
real limits of their knowledge, naturalists of the time hesitated to suggest that certain marine
animals represented by fossils no longer existed anywhere on Earth (Eicher 1976); an organism
might still be living in some deeper, unexplored part of the ocean.
Following the lead of William Smith's work on faunal succession, George Cuvier (1769–1832),
a French zoologist, worked out the stratigraphic sequences of terrestrial vertebrates and marine
invertebrates in the strata of the Paris Basin. In 1812 he showed conclusively that many fossil
vertebrates have no known living counterparts, and people agreed that it was highly unlikely that
such big land animals would go undiscovered (Eicher 1976).
Many geologists consider James Hutton (1726–1797) to be the father of historical geology.
Hutton observed such processes as wave action, erosion by running water, and sediment
transport and concluded that given enough time these processes could account for the geologic
features in his native Scotland. He thought that “the past history of our globe must be explained
by what can be seen to be happening now.” This assumption that present-day processes have
operated throughout geologic time was the basis for the principle of uniformitarianism.
Before Hutton, no one had effectively demonstrated that geologic processes occurred over long
periods of time. Hutton persuasively argued that seemingly weak, slow-acting processes could,
over long spans of time, produce effects that were just as great as those resulting from sudden
catastrophic events. And, unlike his predecessors, Hutton cited verifiable observations to support
his ideas.
Although Hutton developed a comprehensive theory of uniformitarian geology, Charles Lyell
(1797–1875) became its principal advocate. Lyell was successful in interpreting and publicizing
uniformitarianism for society at large. Hutton’s idea of uniformitarianism (and his cumbersome
and difficult literary style) had simply failed to capture the imagination of his generation, so
geologists often credit Lyell with advancing the basic principles of modern geology. Lyell’s
Principles of Geologyis a landmark text in the history of science and as important to modern
world views as the works of Charles Darwin. In 1990 the University of Chicago Press
republished his works. In the first of three volumes, Charles Lyell sets forth the uniformitarian
argument: processes now visibly acting in the natural world are essentially the same as those that
have acted throughout the history of the Earth, and are sufficient to account for all geologic
Like William Smith, Cuvier recognized distinctive fossils, but he emphasized the changes in
them over time. As Cuvier carefully worked out a succession of different faunas in the strata of
the Paris Basin, he noted that the younger deposits contained creatures more like those of the
present day than did the older deposits (Eicher 1976). Therefore, in addition to proving the
extinction of species, Cuvier demonstrated that each changing sequence of faunas represents a
particular time in the geologic past.
Rocks that contain fossils occur in a very real and understandable order. Rocks of certain time
periods can be recognized and separated by their fossil content (Boggs, 2012). This is a skill that
geologists acquire as they do field work and explore the Earth! The fauna from the Mississippian
is very different from the Ordovician and easily distinguishable! Groups of fossils, or fossil
assemblages, can be used to correlate rock units across continents.
Geological Structures: After carefully reading this chapter, completing the exercises within it,
and answering the questions at the end, you should be able to:
● Describe the types of stresses that exist within the Earth’s crust
● Explain how rocks respond to those stresses by brittle, elastic, or plastic deformation, or
by fracturing
● Summarize how rocks become folded and know the terms used to describe the features
of folds
● Describe the conditions under which rocks fracture
● Summarize the different types of faults, including normal, reverse, thrust, and
● Measure the strike and dip of a geological feature
● Plot strike and dip information on a map
12.1 Stress and Strain
Rocks are subject to stress —mostly related to plate tectonics but also to the weight of overlying
rocks—and their response to that stress is strain (deformation). In regions close to where plates
are converging stress is typically compressive—the rocks are being squeezed. Where plates are
diverging the stress is extensive—rocks are being pulled apart. At transform plate boundaries,
where plates are moving side by side there is sideways or shear stress—meaning that there are
forces in opposite directions parallel to a plane. Rocks have highly varying strain responses to
stress because of their different compositions and physical properties, and because temperature is
a big factor and rock temperatures within the crust can vary greatly.
We can describe the stress applied to a rock by breaking it down into three dimensions—all at
right angles to one-another (Figure 12.2). If the rock is subject only to the pressure of burial, the
stresses in all three directions will likely be the same. If it is subject to both burial and tectonic
forces, the pressures will be different in different directions.
Figure 12.2
Depiction of the stress applied to rocks within the crust. The stress can be broken down into three
components. Assuming that we’re looking down in this case, the green arrows represent
north-south stress, the red arrows represent east-west stress, and the blue arrows (the one
underneath is not visible) represent up-down stress. On the left, all of the stress components are
the same. On the right, the north-south stress is least and the up-down stress is greatest. [SE]
Rock can respond to stress in three ways: it can deform elastically, it can deform plastically, and
it can break or fracture. Elastic strain is reversible; if the stress is removed, the rock will return
to its original shape just like a rubber band that is stretched and released. Plastic strain is not
reversible. As already noted, different rocks at different temperatures will behave in different
ways to stress. Higher temperatures lead to more plastic behaviour. Some rocks or sediments are
also more plastic when they are wet. Another factor is the rate at which the stress is applied. If
the stress is applied quickly (for example, because of an extraterrestrial impact or an earthquake),
there will be an increased tendency for the rock to fracture. Some different types of strain
response are illustrated in Figure 12.3.
Figure 12.3
The varying types of response of geological materials to stress. The straight dashed parts are
elastic strain and the curved parts are plastic strain. In each case the X marks where the material
fractures. A, the strongest material, deforms relatively little and breaks at a high stress level. B,
strong but brittle, shows no plastic deformation and breaks after relatively little elastic
deformation. C, the most deformable, breaks only after significant elastic and plastic strain. The
three deformation diagrams on the right show A and C before breaking and B after breaking.
The outcomes of placing rock under stress are highly variable, but they include fracturing, tilting
and folding, stretching and squeezing, and faulting. A fracture is a simple break that does not
involve significant movement of the rock on either side. Fracturing is particularly common in
volcanic rock, which shrinks as it cools. The basalt columns in Figure 12.4a are a good example
of fracture. Beds are sometimes tilted by tectonic forces, as shown in Figure 12.4b, or folded as
shown in Figure 12.1.
Figure 12.4 Rock
structures caused by various types of strain within rocks that have been stressed [all by SE]
When a body of rock is compressed in one direction it is typically extended (or stretched) in
another. This is an important concept because some geological structures only form under
compression, while others only form under tension. Most of the rock in Figure 12.4c is
limestone, which is relatively easily deformed when heated. The dark rock is chert, which
remains brittle. As the limestone stretched (parallel to the hammer handle) the brittle chert was
forced to break into fragments to accommodate the change in shape of the body of rock. A fault
is a rock boundary along which the rocks on either side have been displaced relative to each
other (Figure 12.4d).
12.2 Folding
When a body of rock, especially sedimentary rock, is squeezed from the sides by tectonic forces,
it is likely to fracture and/or become faulted if it is cold and brittle, or become folded if it is
warm enough to behave in a plastic manner.
The nomenclature and geometry of folds are summarized on Figure 12.5. An upward fold is
called an anticline, while a downward fold is called a syncline. In many areas it’s common to
find a series of anticlines and synclines (as in Figure 12.5), although some sequences of rocks are
folded into a single anticline or syncline. A plane drawn through the crest of a fold in a series of
beds is called the axial plane of the fold. The sloping beds on either side of an axial plane are
limbs. An anticline or syncline is described as symmetrical if the angles between each of limb
and the axial plane are generally similar, and asymmetrical if they are not. If the axial plane is
sufficiently tilted that the beds on one side have been tilted past vertical, the fold is known as an
overturned anticline or syncline.
Figure 12.5
Examples of different types of folds and fold nomenclature. Axial planes are only shown for the
anticlines, but synclines also have axial planes. [SE]
A very tight fold, in which the limbs are parallel or nearly parallel to one another is called an
isoclinal fold (Figure 12.6). Isoclinal folds that have been overturned to the extent that their
limbs are nearly horizontal are called recumbent folds.
Figure 12.6
An isoclinal recumbent fold [SE]
Folds can be of any size, and it’s very common to have smaller folds within larger folds (Figure
12.7). Large folds can have wavelengths of tens of kilometres, and very small ones might be
visible only under a microscope. Anticlines are not necessarily, or even typically, expressed as
ridges in the terrain, nor synclines as valleys. Folded rocks get eroded just like all other rocks
and the topography that results is typically controlled mostly by the resistance of different layers
to erosion (Figure 12.8).Figure 12.7 Folded limestone (grey) and chert (rust-coloured) in Triassic
Quatsino Formation rocks on Quadra Island, B.C. The image is about 1 m across. [SE]
Figure 12.8
Example of the topography in an area of folded rocks that has been eroded. In this case the green
and grey rocks are most resistant to erosion, and are represented by hills. [SE]
12.3 Fracturing and Faulting
A body of rock that is brittle—either because it is cold or because of its composition, or both—
is likely to break rather than fold when subjected to stress, and the result is fracturing or faulting.
Fracturing is common in rocks near the surface, either in volcanic rocks that have shrunk on
cooling (Figure 12.4a), or in other rocks that have been exposed by erosion and have expanded
(Figure 12.9).
Figure 12.9 Granite in the
Coquihalla Creek area, B.C. (left) and sandstone at Nanoose, B.C. (right), both showing
fracturing that has resulted from expansion due to removal of overlying rock. [SE]
A fracture in a rock is also called a joint. There is no side-to-side movement of the rock on either
side of a joint. Most joints form where a body of rock is expanding because of reduced pressure,
as shown by the two examples in Figure 12.9, or where the rock itself is contracting but the body
of rock remains the same size (the cooling volcanic rock in Figure 12.4a). In all of these cases,
the pressure regime is one of tension as opposed to compression. Joints can also develop where
rock is being folded because, while folding typically happens during compression, there may be
some parts of the fold that are in tension (Figure 12.10).
Figure 12.10
A depiction of joints developed in the hinge area of folded rocks. Note that in this situation some
rock types are more likely to fracture than others. [SE]
Finally joints can also develop when rock is under compression as shown on Figure 12.11, where
there is differential stress on the rock, and joint sets develop at angles to the compression
Figure 12.11 A depiction of joints developed in a rock that is under stress. [SE]
A fault is boundary between two bodies of rock along which there has been relative motion
(Figure 12.4d). As we discussed in Chapter 11, an earthquake involves the sliding of one body of
rock past another. Earthquakes don’t necessarily happen on existing faults, but once an
earthquake takes place a fault will exist in the rock at that location. Some large faults, like the
San Andreas Fault in California or the Tintina Fault, which extends from northern B.C. through
central Yukon and into Alaska, show evidence of hundreds of kilometres of motion, while others
show less than a millimetre. In order to estimate the amount of motion on a fault, we need to find
some geological feature that shows up on both sides and has been offset (Figure 12.12).
12.12 A
fault (white
line) in
rocks on
Island, B.C.
The pink
dyke has
been offset
by the fault
and the
extent of the offset is shown by the white arrow (approximately 10 cm). Because the far side of
the fault has moved to the right, this is a right-lateral fault. If the photo were taken from the
other side, the fault would still appear to have a right-lateral offset. [SE]
There are several kinds of faults, as illustrated on Figure 12.13, and they develop under different
stress conditions. The terms hanging wall and footwall in the diagrams apply to situations where
the fault is not vertical. The body of rock above the fault is called the hanging wall, and the body
of rock below it is called the footwall. If the fault develops in a situation of compression, then it
will be a reverse fault because the compression causes the hanging wall to be pushed up relative
to the footwall. If the fault develops in a situation of extension, then it will be a normal fault,
because the extension allows the hanging wall to slide down relative to the footwall in response
to gravity.
The third situation is where the bodies of rock are sliding sideways with respect to each other, as
is the case along a transform fault (see Chapter 10). This is known as a strike-slip faultbecause
the displacement is along the “strike” or the length of the fault. On strike-slip faults the motion is
typically only horizontal, or with a very small vertical component, and as discussed above the
sense of motion can be right lateral (the far side moves to the right), as in Figures 12.12 and
12.13, or it can be left lateral (the far side moves to the left). Transform faults are strike-slip
Figure 12.13
Depiction of reverse, normal, and strike-slip faults. Reverse faults happen during compression
while normal faults happen during extension. Most strike-slip faults are related to transform
boundaries. [SE after:
In areas that are characterized by extensional tectonics, it is not uncommon for a part of the
upper crust to subside with respect to neighbouring parts. This is typical along areas of
continental rifting, such as the Great Rift Valley of East Africa or in parts of Iceland, but it is also
seen elsewhere. In such situations a down-dropped block is known as a graben (German for
ditch), while an adjacent block that doesn’t subside is called a horst (German for heap) (Figure
12.14). There are many horsts and grabens in the Basin and Range area of the western United
States, especially in Nevada. Part of the Fraser Valley region of B.C., in the area around Sumas
Prairie is a graben.
Figure 12.14
Depiction of graben and horst structures that form in extensional situations. All of the faults are
normal faults. [SE]
A special type of reverse fault, with a very low-angle fault plane, is known as a thrust fault.
Thrust faults are relatively common in areas where fold-belt mountains have been created during
continent-continent collision. Some represent tens of kilometres of thrusting, where thick sheets
of sedimentary rock have been pushed up and over top of other rock (Figure 12.15).
Figure 12.15 Depiction a thrust
fault. Top: prior to faulting. Bottom: after significant fault offset. [SE]
There are numerous thrust faults in the Rocky Mountains, and a well-known example is the
McConnell Thrust, along which a sequence of sedimentary rocks about 800 m thick has been
pushed for about 40 km from west to east (Figure 12.16). The thrusted rocks range in age from
Cambrian to Cretaceous, so in the area around Mt. Yamnuska Cambrian-aged rock (around 500
Ma) has been thrust over, and now lies on top of Cretaceous-aged rock (around 75 Ma) (Figure
Figure 12.16
Depiction of the McConnell Thrust in the eastern part of the Rockies. The rock within the faded
area has been eroded. [SE]
Figure 12.17
The McConnell Thrust at Mt. Yamnuska near Exshaw, Alberta. Carbonate rocks (limestone) of
Cambrian age have been thrust over top of Cretaceous mudstone. [SE]
12.4 Measuring Geological Structures
Geologists take great pains to measure and record geological structures because they are
critically important to understanding the geological history of a region. One of the key features to
measure is the orientation, or attitude, of bedding. We know that sedimentary beds are deposited
in horizontal layers, so if the layers are no longer horizontal, then we can infer that they have
been affected by tectonic forces and have become either tilted, or folded. We can express the
orientation of a bed (or any other planar feature) with two values: first, the compass orientation
of a horizontal line on the surface—the strike—and second, the angle at which the surface dips
from the horizontal, (perpendicular to the strike)—the dip (Figure 12.18).
It may help to imagine a vertical surface, such as a wall in your house. The strike is the compass
orientation of the wall and the dip is 90˚ from horizontal. If you could push the wall so it’s
leaning over, but still attached to the floor, the strike direction would be the same, but the dip
angle would be less than 90˚. If you pushed the wall over completely so it was lying on the floor,
it would no longer have a strike direction and its dip would be 0˚. When describing the dip it is
important to include the direction. In other words. if the strike is 0˚ (i.e., north) and the dip is 30˚,
it would be necessary to say “to the west” or “to the east.” Similarly if the strike is 45˚ (i.e.,
northeast) and the dip is 60˚, it would be necessary to say “to the northwest” or “to the
Measurement of geological features is done with a special compass that has a built-in clinometer,
which is a device for measuring vertical angles. An example of how this is done is shown on
Figure 12.19.
Figure 12.18
A depiction of the strike and dip of some tilted sedimentary beds partially covered with water.
The notation for expressing strike and dip on a map is shown. [SE]
Figure 12.19 Measurement of strike (left) and dip (right) using a geological compass with a
clinometer. [SE]
Strike and dip are also used to describe any other planar features, including joints, faults, dykes,
sills, and even the foliation planes in metamorphic rocks. Figure 12.20 shows an example of how
we would depict the beds that make up an anticline on a map.
Figure 12.20
A depiction of an anticline and a dyke in cross-section (looking from the side) and in map view
(a.k.a. plan view) with the appropriate strike-dip and anticline symbols. [SE]
The beds on the west (left) side of the map are dipping at various angles to the west. The beds on
the east side are dipping to the east. The middle bed (light grey) is horizontal; this is denoted by a
cross within a circle. The dyke is dipping at 80˚ to the west. The hinge of the fold is denoted with
a dashed line with two arrows point away from it. If it were a syncline, the arrows would point
towards the line.
The topics covered in this chapter can be summarized as follows:
Stress within rocks, which includes compression, extension and
shearing, typically originates from plate-boundary processes. Rock that
Stress and is stressed responds with either elastic or plastic strain, and will
eventually break. The way a rock responds to stress depends on its
composition and structure, the rate at which strain is applied, and also
to the temperature of the rock body and the presence of water.
Folding is generally a plastic response to compressive stress, although
some brittle behaviour can happen during folding. An upward fold is
an anticline. A downward fold is a syncline. The axis of a fold can be
vertical, inclined, or even horizontal. If we know that the folded beds
have not been overturned, then we can use the more specific terms:
anticline and syncline.
Fractures (joints) typically form during extension, but can also form
Fracturing during compression. Faulting, which involves the displacement of
rock, can take place during compression or extension, as well as during
shearing at transform boundaries. Thrust faulting is a special form of
reverse faulting.
It is important to be able to measure the strike and dip of planar
surfaces, such as a bedding planes, fractures or faults. Special symbols
are used to show the orientation of structural features on geological
Geologic Structures
Types of geologic structures:
(1) Primary structures: those which develop at the time of formation of the rocks (e.g.
sedimentary structures, some volcanic structures, .... etc.).
(2) Secondary structures: which are those that develop in rocks after their formation as a result
of their subjection to external forces.
(3) Compound structures: form by a combination of events some of which are contemporaneous
with the formation of a group of rocks taking part in these "structures".
Stress: is the force applied over a given area of the rock mass. It is of three different kinds:
(1) Compressional stress which tends to squeeze the rock
(2) Tensional stress, which tends to pull a rock apart
(3) Shear stress, which results from parallel forces that act on different parts of the rock
body in opposite directions.
Strain: Is the change in the shape or size of a rock in response to stress. A rock is said to deform
elastically if it can return to its original size once the stress is removed. Plastic deformation on
the other hand, results in permanent changes in the size and shape of the rock, even after the
stress is removed. Plastic deformation of a rock is also known as ductile deformation. After
deforming plastically for some time, a rock which continues to be subjected to stress may finally
break, a behaviour known as brittle deformation.
Factors affecting how a rock deforms:
Depth: Lithostatic pressure + heat
Therefore, a rock may undergo ductile deformation when subjected to stress at certain depths
within the earth where pressures and temperatures are relatively high, or if fluids are abundant,
but the same rock may undergo brittle deformation at shallower depths.
Measuring geological structures:
Strike: (direction)
Dip: (direction & angle)
A- Secondary structures
Types of secondary geologic structures:
(a) folds, which are a form of ductile deformation, and (b) fractures, represented by faults and
joints which generally result from the brittle behaviour of rocks in response to stress.
I- Folds
Folds are bends or flexures in the earth's crust, and can therefore be identified by a change in the
amount and/or direction of dip of rock units. Most folds result from the ductile deformation of
rocks when subjected to compressional or shear stress. In order to understand and classify folds,
we must study their forms and shapes, and be able to describe them. The following definitions
are therefore essential for the description of a fold:
1- Hinge line: Is the line of maximum curvature on a folded surface. The hinge line almost
always coincides with the axis of the fold defined as a line lying in the plane that bisects a fold
into two equal parts.
2- The axial plane is an imaginary plane dividing the fold into two equal parts known as limbs.
It is therefore the plane which includes all hinge lines for different beds affected by the same
3- The crest of a fold can be considered the highest point on a folded surface. The trough is the
lowest point on a folded surface.
4- The interlimb angle: Is the angle between two limbs of the same fold. It is measured in a
plane perpendicular to that of the fold axis.
5- The angle of plunge of a fold is the angle between the fold axis and the horizontal plane,
measured in a vertical plane. The direction of plunge of a fold is the direction in which the fold
axis dips into the ground from the horizontal plane.
6- The median surface: Is the surface that passes through points where the fold limb changes its
curvature from concave to convex.
7- The amplitude of a fold: is the vertical distance between the median surface and the fold
hinge, both taken on the same surface of the same folded unit.
8- The wavelength of a fold system is the distance between two consecutive crests or troughs
taken on the same folded surface.
Classification of folds
Folds may be classified based on the direction of dip of their limbs, the inclination of their axial
planes, the value of their interlimb angle, their plunge, and their general shape and effects on the
thickness of the folded layers. In order to describe a fold correctly, one may have to use more
that one of these classifications; e.g. recumbent anticline, open syncline, tight plunging anticline,
.... etc. (see below).
(a) Classification based on the direction of dip of the limbs:
When both limbs of a fold dip away from the fold axis, the fold is called an antiform. If both
limbs dip towards the fold axis, the fold is known as a synform. If the relative ages of the folded
units are known, such that the oldest units occur in the core of the antiform, the antiform is called
"anticline". Similarly, if the youngest units occur in the "center" of a synformal structure, it is
known as a syncline (Fig. 1).
A monocline is a single step-like bend in a rock unit, and is often caused by vertical
displacement. A dome consists of uparched rocks that dip in all directions away from the central
point. A basin is a downwarp in which the layers dip in all directions from all sides towards the
centre (Fig. 2). A fold is described as isoclinal if both limbs dip in the same direction at the same
angle (Fig. 3).
(b) Classification based on the inclination of the axial plane: (Fig. 4)
A symmetrical (or upright) fold is one in which the axial plane bisects the fold (and is vertical).
If the axial plane is inclined at an angle < 45° (measured from the vertical plane), the fold is said
to be inclined. If the angle of inclination of the axial plane is > 45° (from the vertical plane),
then both limbs of the fold will dip in the same direction, and the fold is known as inverted or
overturned. If the axial plane is horizontal, the fold is known as recumbent.
(c) Classification based on the value of the interlimb angle (Fig. 5):
(1) Open folds: those with an interlimb angle > 70°, (2) Closed folds: with interlimb angles
between 30 and 70°, (3) Tight folds: with interlimb angles < 30°, (4) Isoclinal folds: have zero
interlimb angles.
II- Faults
A fault is a fracture in the earth's rock units along which there has been an observable amount of
movement and displacement. Unlike folds which form predominantly by compressional stress,
faults result from either tension, compression or shear.
In order to correctly describe a fault, it is essential to understand its components:
1- The fault plane: Is the plane of dislocation or fracture along which displacement has
occurred. The fault plane therefore separates one or more rock units into two blocks.
2- The Hanging wall and footwall blocks: If the fault plane is not vertical, then the block lying
on top of the fault plane is known as the hanging wall block, whereas that lying below this plane
is known as the footwall block.
3- The downthrown and upthrown blocks: The downthrown block is the one that has moved
downwards relative to the other block, whereas the upthrown block is that which registers an
upward relative movement.
4- The Dip of the fault plane is the angle of inclination of the fault plane measured from the
horizontal plane perpendicular to its strike.
6- Fault Throw: Is the vertical displacement of a fault.
8- Dip slip: Is the amount of displacement measured on the fault plane in the direction of its dip.
9- Strike slip: Is the amount of displacement measured on the fault plane in the direction of its
10- Net slip: Is the total amount of displacement measured on the fault plane in the direction of
N.B. In measuring the slip or throw of a fault, the displacement has to be measured using the
same surface of the same unit affected by that fault.
Types of Faults
1- Normal fault: Is a fault in which the hanging wall appears to have moved downwards relative
to the footwall (i.e. downthrown block = hanging wall block).
2- Reverse fault: Is a fault in which the hanging wall appears to have moved upwards relative to
the footwall (i.e. upthrown block = hanging wall block). Because the displacement in both
normal and reverse faults occurs along the dip of the fault plane, they may be considered types of
dip slip faults.
3- Thrust fault (or thrust): Is a reverse fault in which the fault plane is dipping at low angles (<
45°). Thrusts are very common in mountain chains (fold and thrust belts) where they are
characterized by transporting older rocks on top of younger ones over long distances.
4- Strike slip (wrench, tear or transcurrent) fault: Is a fault in which the movement is
horizontal along the strike of the fault plane. Strike slip faults are either dextral or sinistral.
When viewed on end (Fig. 13), a dextral fault (also known as right lateral fault) is one in which
the block on the observer's right hand side appears to have moved towards him, whereas a
sinistral strike slip fault (also known as left lateral) is one in which the block on the observer's
left hand side appears to have moved towards him.
5- Oblique slip fault: is one in which the displacement was both in the strike and dip directions
(i.e. the displacement has strike and dip components). Keep in mind that an oblique slip fault can
also be either normal or reverse.
From this classification of faults, it can be seen that normal faults result predominantly from
tensional stress, reverse faults and thrusts from compression (or shear), and strike slip faults from
tension, compression or shear.
Fault Associations and Fault Systems (Fig. 6)
Faults often occur in groups. If two normal faults have parallel strikes and share the same
downthrown block, a trough-like structure results which is known as a graben. A horst is an
uplifted block bounded by two normal faults that strike parallel to each other (and which share
the same upthrown block Þ the horst). Grabens and horsts are common in areas of very early
rifting (e.g. the East African Rift Valley). Step faults are several faults with parallel strikes and a
repeated downthrow in the same direction giving the area an overall step - like appearance. They
are common in rifted areas (e.g. on the flanks of the Red sea).
Geomorphological features associated with faults:
Fault planes often result in the exposure of units that erode easily along the fault trace resulting
in the development of valleys or the control of stream flow. In other cases, faults cause the offset
of streams, causing them to bend sharply when they intersect the fault plane. The topography
may also be strongly influenced by faulting so that the fault plane can be identified on the ground
by a sudden and sharp change in elevation, known as a fault scarp.
Recognition of movement along fault planes
Movement along a fault plane can often be recognized by the following criteria:
1- Fault drag: where small - scale folding or warping of units takes place as a result of the
dragging forces along the fault plane (Fig. 7).
2- Fault breccia and fault gouge: As a result of movement along the fault plane, rocks are often
broken up into sharp angular pieces known as breccia. The fragments may be further crushed
into powder - like material, known as fault gouge.
3- Slickensides: As a result of movement and friction along the fault plane, this plane may
become highly polished or abraded with striations that are known as slickensides (Fig. 8).
III- Joints
Joints are fractures in the rocks characterized by no movement along their surfaces. Although
most joints are secondary structures, some are primary, forming at the time of formation of the
Types of joints
1- Columnar joints: Are joints that form in basalts. When the basaltic lava cools, it contracts
giving rise to hexagonal shaped columns.
2- Mud cracks: Are joints that form in mud. As the mud loses its water, it contracts and cracks.
3- Secondary joints: Are joints that form in rocks as a result of their subjection to any form of
stress (compression, tension or shear). Joints that are oriented in one direction approximately
parallel to one another make up a joint set. Rocks often have more than one set of joints with
different orientations, which may intersect, and are then known as joint systems (Fig. 9). Note
that tensional stress usually results in one set of joints, whereas compression may form more than
one set.
4- Sheet joints: Are joints that form in granitic rocks in deserts causing them to break into thin
parallel sheets. These joints form when the rocks expand as a result of the rapid removal of the
overlying rock cover, possibly due to faulting or quarrying. This process is called exfoliation.
B- Compound Structures
An unconformity is a surface (or contact) along which there was no fracturing (i.e. not a fault or
joint) and which represents a break in the geologic record. An unconformity therefore indicates a
lack of continuity of sedimentary deposition in an area, resulting in rocks of widely different
ages occurring in contact with each other. Unconformities usually result from changes in the
sedimentary history of an area, which may be due to vertical movements (e.g. uplift followed by
erosion and deposition), deformation (also followed by deposition), changes in sea level (which
may be due to climatic changes, among other things), ...etc.
In many cases, unconformities represent a buried erosional surface. In such cases, erosion of the
older units results in their fragmentation into smaller pieces. As soon as deposition resumes,
these fragments may consolidate to form a rock known as breccia (if the fragments are angular)
or conglomerate (if the fragments are rounded). Because the breccia or conglomerate occur at the
base of the younger units lying on top of the unconformity surface, and because their fragments
are derived from the units below this surface, the conglomerates or breccias are known as basal
conglomerates or basal breccias.
Types of unconformities (Fig. 10)
1- Angular unconformities: are those in which the angle of dip of the younger layers is
different from that of the older ones.
2- Disconformities: are those in which the units above and below the unconformity surface are
parallel to each other, but not continuous in deposition or age.
3- Nonconformities: are those in which plutonic or metamorphic rocks are covered by
sedimentary or volcanic units.
Geological Structures and Plate tectonics:
All three types of plate boundaries are characterized by certain deformational (structural)
features. The most intense deformation occurs in areas of continent - continent collision.
Divergent boundaries: Mostly extensional structures; horsts, grabens, step faults, ... etc.
Convergent boundaries: "Fold and thrust belts"; nappes (Fig. 11).
Transform boundaries: Strike slip faults, en echelon fault
Mountain building (Orogeny):
Examples:Andes, Rockies, Himalayas, Alps, Appalachians, Basin & Range: Tetons,
Vertical movements (Epeirogeny/ uplift) and Isostasy.
Epeirogeny is the vertical movement of crustal blocks relative to sea level in a "non-mountain
building" event. The term is rather vague and is almost obsolete! The simple "non-genetic" term
uplift is more useful. Isostasy is the state of balance between extensive blocks of the earth's
crust which rise to different levels and appear at the surface as mountain ranges, plateaus and
plains. Applied to mountains (where the lighter continental crust overlying the denser mantle is
quite thick), this concept dictates that the higher the mountain, the thicker the crust beneath it, or
the deeper the "crustal root" of that mountain within the underlying mantle. Because the sialic
material is less dense than sima, the gravitational attraction beneath mountain chains is much
lower than that on the ocean floors (Fig. 12)
Another important effect of isostasy is seen when material is eroded from a mountain, resulting
in the rise of the crust - mantle boundary (the Moho) to compensate for the eroded material. This
process is known as isostatic readjustment or isostatic rebound (Fig. 13). A good example of this
process can be seen in parts of the Baltics, Arctic and the Great Lakes Region of North America.
During the ice ages, these areas were covered by large ice caps which depressed the crust by as
much as 200 to 300 meters. When the ice melted, the crust rebounded slowly, as evidenced by
the occurrence of beach deposits at high elevations.
Folds jpb, 2017 207 FOLDS The term fold is used when one or stacks of originally flat and
planar surfaces such as sedimentary beds become bent or curved as a result of plastic (i.e.
permanent) and ductile deformation. Folds in rocks vary in size from microscopic crinkles to
mountain-size folds. They occur singly as isolated folds and in extensive fold trains at all scales.
A set of folds distributed on a regional scale constitutes a fold belt. Fold belts are typically
associated with convergent plate boundaries and directed compressive stress. Folds form under
varied conditions of stress, hydrostatic pressure, pore pressure, and temperature as evidenced by
their presence in sediments, sedimentary rocks, the full spectrum of metamorphic rocks, and in
some igneous rocks. Folds may result from a primary deformation, which means that folding
occurred during the formation of the rock, or a consequence of a secondary, i.e. tectonic
deformation. Slumps in soft sediments and flow folds in lavas are examples of primary folds. In
structural geology we are mainly concerned by the tectonic folds that are, in general, produced
by a shortening component parallel to the layering of the rocks. Their spectacular presence in
shear zones and mountain belts indicates that distributed, ductile deformation has resulted in
gradual and continuous changes in a rock layer both in its attitudes and internally. However, it
must be remembered that the absence of folds does not demonstrate the absence of pervasive
deformation. This lecture deals with the description and morphological classification of folds.
Such classifications aim at using the geometry and degree of evolution of folds as indicators of
the amount of deformation, hence displacement and strain patterns involved in the development
of deformed areas. It is still difficult to extract dynamic or kinematic information from these
structures since their shapes in rocks are highly variable. This variability reflects differences in
several types of parameters such as layer thickness, initial layer irregularities, strain intensity,
deformation path, rheology and degree of anisotropy. Folded single surface - basic geometrical
definitions A fold may bend a single surface, or affect one layer bounded by two surfaces or
deform a stack of layers with several interfaces. Since there is considerable variation in fold
morphology a proliferation of terms has developed to describe qualitatively these structures. For
convenience, it is first easier to define folds on a single surface. Morphology of a folded surface:
Hinge, limb, inflections The radius of curvature of a folded surface varies progressively from
point to point. The point of smallest radius of curvature is called the hinge. It is flanked by two
areas of larger radius of curvature: the limbs. Folds are three-dimensional structures. Connecting
the hinge points on a specific folded surface defines the hinge line or fold axis. The fold axis is
the most important structural element of a fold Folds jpb, 2017 208 because it shows the
direction of maximum continuity of this fold. Some folds (e.g. box folds) may have several hinge
lines. The inflection points are points of zero curvature, where the sense of curvature changes
from a convex to a concave line. They usually are aligned on either limb of a fold; The limb may
be then redefined as the fold segment between a hinge line and the adjacent inflection line, which
is the locus of inflection points. If the limb has a straight segment, its midpoint is conveniently
taken as inflection point. Antiform and synform A convex-upward fold is an antiform; a
convex-downward fold is a synform. They often come in pairs. The region towards the inner,
concave side of a folded layer is the core of the fold. The convex and concave sides of a folded
layer are also distinguished as the extrados and the intrados, respectively. Anti- and synclinorium
are large (megascopic, regional-scale) anti- and synforms, respectively. An oval-shaped antiform
with no distinct trend of the hinge line, in which layering dips outward from a central point, is
termed a dome, a synform with no distinct trend of hinge line, i.e. in which layering dips inward
toward a central point, is a basin. Folds that are neither antiformal nor synformal, whose limbs
converge sidewise, are called neutral folds. They comprise vertically plunging folds, folds-with
horizontal axial surfaces, and folds that plunge parallel to the dips of their axial surface. Folds
jpb, 2017 209 Anticline and syncline Anticline and syncline are terms with stratigraphic
significance. Anticlines have older strata in the core. Synclines have younger strata at the core. In
simply-folded areas, anticlines are antiformal and synclines are synformal. But antiformal
synclines and synformal anticlines may exist in refolded regions. In these regions it is important
to determine the younging direction, which is the direction in which the strata become younger
along the axial surface. Profile A folded surface is fully described in a plane perpendicular to the
fold axis. The transverse profile (or simply profile) of a fold is the section drawn perpendicular
to the fold axis and its axial surface; this contrasts with a geological section which is normally
drawn in a vertical plane. Folds jpb, 2017 210 The profile is a reference plane used to describe
and measure all geometrical characteristics of the fold: height or amplitude, wavelength,
tightness, roundness. Indeed, these aspects vary with the angular relationship between any
section plane and the folded surface. Together, all various geometrical features of the fold profile
and the orientation of the fold axis define the style of a fold. Crest, culmination and trough The
crest is the highest point of an antiform, the lowest point of a synform is the trough. Imaginary
lines joining crest and trough points of any bedding surface are crest lines or trough lines. They
are also the lines where layering changes orientation from dip in one direction to dip in the
opposite direction, away from each other along the crest line, toward one another along the
trough line. Crest and trough lines are neither horizontal nor rectilinear but vary in height along
their length. The high points in crest lines are referred to as culminations and the low points in
trough lines are depressions. Folds jpb, 2017 211 Interlimb angles In profile, the smaller angle
made by the limbs of a fold is termed the inter-limb angle, a measure of the tightness of the fold.
It is the angle subtended by the tangents at two adjacent inflection points, which may reflect the
intensity of compression. A qualitative classification based on the interlimb angle, separates five
tightness classes: inter-limb angle tightness class 180 to ca. 120° Gentle 120 -- 70° Open 70 -30° Close less than 30° Tight 0°, i.e. parallel limbs Isoclinal < 0° Fan Folds jpb, 2017 212 Fan
folds have negative interlimb angles. A cusp is a fold where both hinge and inflection points are
the same point; in other words the fold has no inflection point. Its tightness is defined by a
cusp-angle between the tangents to the folded surface at the cusp. Fold closures The fold closure
indicates the direction in which the limbs converge. For example, a fold closes westward. The
shape of the fold closure depends on how the curvature of the folded surface changes around the
hinge; it may be very sharp and the limbs relatively straight, or the curvature more regular
around the fold. Fold closures are thus qualitatively described as rounded or angular. Arrowhead
folds or flame folds have sharp hinges with distinctly, often sigmoidally curved limbs. Kinks are
folds with straight, planar limbs (there is no inflexion point) and angular hinges (the hinge zone
is reduced to a point). They form in strongly anisotropic rocks in which the well-developed
anisotropy is either thin laminated beds or foliation planes. Where kinks are markedly
asymmetrical, the long narrow zone defined by the tabular, short limb is referred to as a kink
band and the axial plane traces are referred to as kink band boundaries. Folds jpb, 2017 213 The
bluntness ratio is a quantitative measure of how round or angular the hinge is. It is defined as
B=r r h i where hr is the radius of curvature at the hinge and i r the radius of the circle tangent to
the limbs at the two inflection points. The angle θ between the two i r is sometimes used to
define the tightness of the fold. Cylindricity Cylindrical folds Folds are often drawn as
cylindrical structures, meaning that the fold axis is a straight line which, when moved parallel to
itself, generates any single fold of the same generation. The axis of cylindricity is parallel to the
fold axis. In three dimensions, a cylindrical fold appears as a straight line in a section parallel to
its axis, whereas in any other section the trace of the folded surface has a wavy shape. Doubly
plunging folds However, hinge lines are rarely straight. Non-cylindrical folds deviate from the
ideal cylindrical geometry. Hinges of non-cylindrical folds are curved within a plane
(curvilinear) and, therefore, change in trend and plunge. A conical fold describes a
non-cylindrically folded surface that has the approximate geometry of a cone. The plunge of the
hinge line reverses along a doubly plunging fold. If the hinge line plunges away from a high
point (the axis is convex upward), the high point is a Folds jpb, 2017 214 culmination; if it
plunges toward a low point (the axis is concave upward) the low point is a depression. Nearly
circular culminations and depressions are domes and basins, respectively. Sheath folds A sheath
fold has a strongly curved hinge line sweeping around through an arc of more than 90° up to
hairpin bend. Sheath folds contain a long (stretching) axis along the length of the tube or tongue
shape, whilst cross sections normal to this axis display closed geometries. Such elliptical sections
or nested rings define eye-folds. These “tubular” folds generally reflect heterogeneous simple
shear or flow superimposed on very simple buckles or perturbation of the simple shear flow such
as, for example a foliation bulge around a rigid clast. During fold amplification due to
progressive shear, the fold axes may behave passively and rotate towards the shear direction until
at high strain they become sub-parallel to the shear direction. Gentle bends of the initial buckle
hinges are therefore accentuated during subsequent shearing to create folds with strongly curved
axes. Sheath folds are therefore characteristic of strongly deformed parts of shear zones.
Buckling of an interface Buckle folding may also affect the planar interface between materials of
contrasting viscosity. When this occurs, the folds have a characteristic form; those closing in one
direction have a broad rounded shape (lobate folds) and hinges closing in the opposite direction
have a narrow or cuspate shape. The cusps always point towards the material with the higher
viscosity. Thus, in outcrops dominated by cuspate-lobate forms, it is possible to know at a glance
which layers was relatively stiffer compared to adjacent beds at the time of folding. This is a
common cause of mullion structures. Folds jpb, 2017 215 Homocline - Monocline A succession
of beds with uniform parallel attitudes over a large area forms a homocline. An antiform and an
adjacent synform delimit a single limb. Such a flexure pair involving a local increase in regional
dip (i.e. only one tilted, step-like limb in an otherwise subhorizontal or gently dipping sequence)
constitutes a monocline. Conversely, a local decrease in regional dip is a structural terrace.
Monoclines and structural terraces are typically large-scale structures along margins of broad
basins or uplift platforms in cratons. Orientation of a plane By definition, folding implies some
rotation of the layering and this rotation may reach large finite values (90° or more). As a
consequence, folded areas are characterised by beds that are no longer horizontal, but instead
inclined. First we must know how to measure and record the inclinations and their directions.
The attitude is the orientation in space of any structure, which requires several measurements.
Geometrically, we know that two intersecting lines define one single plane. - The strike is the
direction of the horizontal line within a sloping bed of rock (i.e. the intersection of an inclined
geological plane with an imaginary horizontal surface). - The direction of dip, the line of
maximum slope on the bed, is perpendicular to the strike. - The dip is the angle between the bed
and the horizontal. It is measured in the imaginary vertical plane orthogonal to the strike (i.e. that
contains the dip direction). Folds jpb, 2017 216 The azimuth of the strike direction, the dip and
the dip direction wholly define any geological plane. Fold systems and folded multilayers - more
definitions A fold may bend a single surface, or affect one layer bounded by two surfaces or
deform a stack of layers with several interfaces. The two sides of a folded layer are also
distinguished with respect to the fold core as the outer arc and inner arc, respectively. A fold
system is a group of folds, often of variable shape, size and orientation, yet spatially and
genetically related. Much of the geometry of folds is concerned with the shape of the profile. On
the transverse profile, the wavy trace of the folded surface may be represented by a function y =
f(x), where the x-axis joins two consecutive inflection points and the y-axis (with positive y
directed upward) is at a right angle to the x-axis. In this case, the inflection points are those
where 2 2 d y dx 0 = . At a hinge point the absolute value of 2 2 d y dx will be maximal.
Orientation of folds A fold, in general, affects several superposed layers. The imaginary surface
connecting the hinge lines on successive layer surfaces of the same fold is the axial plane. This
imaginary plane is often curved and equidistant from each limb and thus often bisects the angle
between them. The orientation of folds is completely given by the fold axis treated as a line and
the attitude of the axial plane (strike and dip), together. Folds jpb, 2017 217 A general rule is that
both the trend and plunge of minor order folds can be used for extrapolation in fieldwork. They
also indicate the trend and plunge of first order folds of the same generation. Fold axis
orientation The orientation of a fold axis is expressed by its plunge and its plunge azimuth: - The
plunge is the inclination measured from the horizontal in the imaginary vertical plane containing
the line). - The direction of plunge (the trend) is the strike (azimuth, the bearing relative to
North) of the imaginary vertical plane that contains the line and the direction in which the
downward inclination occurs. Axial plane orientation The axial planes cuts the hinge zone of a
folded surface along the fold axis. The orientation of the axial plane is expressed as a dip and
strike or, in a more compact form, as a dip and direction of dip. As for any surface the strike is
the trend of the horizontal line contained in the surface; the angle dip is the angle between the
surface and the horizontal plane. The direction of dip of a surface is the trend of the line
perpendicular to the strike of the surface looking down the dip. A semi-quantitative classification
is valid for folds with subhorizontal axes. Upright folds have approximately vertical axial
surfaces. Folds with dipping axial planes are inclined (80° < steeply < 60° < moderately < 30° <
gently < 10°) if the steeper limb has an upward-younging to vertical stratigraphy. The fold is an
overfold if Folds jpb, 2017 218 both limbs dip in the same direction as the axial plane and the
steep limb has an up-side-down stratigraphy. Folds with sub-horizontal axial planes are
recumbent. Fold nappes are large recumbent folds with inverted limb over several kilometers.
Plunging folds have axial planes rotated by more than 90°. Polyclinal folds belong to groups of
folds with sub-parallel hinge lines but non-parallel axial surfaces. Layer thickness - parallel and
similar folds Changes in layer thicknesses reflect material properties of the fold. There are two
end-member geometries: parallel and similar folds. Approximations to these two morphologies
are found in rocks, but the majority of natural folds lies somewhere between the two.
Furthermore, some layers may approximate one of the ideal morphologies while other layers do
not. Parallel folds A fold is parallel if the layer thickness, measured normal to the bed, is constant
all around the fold. In other words, the strata are bent in parallel curves. There are two types: Rounded forms have smoothly curved limbs and broad hinges. - Angular forms have straight
limbs and narrow hinge zones. Rounded, parallel folds In profiles of concentric folds, the folded
surfaces define circular arcs with a common centre. This geometry generates a space problem.
Rounded antiforms reduce downward to a point (cusp). Similarly, deep and wide synforms
wedge out upward. This geometrical limitation requires that parallel folds die out at depth.
Incompetent beds below the antiformal cusps are squeezed in the anticlines. The lower boundary
of the incompetent levels remains essentially unfolded. The incompetent layer between the
apparently undisturbed, flat footwall and the independently folded hanging wall is a
“décollement” horizon. Folds jpb, 2017 219 Ptygmatic folds (from πτύσσω, to buckle in ancient
Greek) involve an irregularly folded, isolated “layer”, typically a quartzo-feldspathic vein in a
much more ductile schistose or gneissic matrix. They occur in high-grade rocks, mostly
migmatites as trains of rounded and near-parallel, commonly concentric folds in which the
amplitude is large (>10) and the wavelength small with respect to the almost constant layer/vein
thickness (meander-like pattern). They have a lobate, tortuous to squiggled appearance (for
example, limbs fold back on themselves and the interlimb angle is negative) and tend to be
polyclinal; however, they have no axial plane foliation. Convolute folds have markedly
curviplanar axial surfaces and are generally disharmonic (adjacent layers do not have the same
wavelength and amplitude). Like ptygmatic folds, they look like complexly contorted but regular
structures without axial plane cleavage. They are characteristic of slumped soft sediments.
Angular parallel folds Chevron folds are symmetric or slightly asymmetric folds with straight
limbs, sharp angular hinges and often acute interlimb angles. They are common in multilayers of
alternating competent and incompetent layers and thus combine both similar (in incompetent
layers) and parallel (in competent layers) fold geometries. Asymmetrical chevron folds are also
termed zigzag folds. Similar folds Folds in which the layer thickness, measured parallel to the
axial surface, is constant are similar. Similar folds tend to have persistent in profiles, that is, the
folded outline including wavelength, symmetry and general shape of a given layer is repeated by
all adjacent layers: the strata are bent into Folds jpb, 2017 220 similar curves. As a geometrical
consequence, beds do not retain their original thickness throughout and the limbs are thinner than
the hinges. Such folds do not die out upward or downward, but maintain the same curvature in
the hinges. This is the case for kink folds, in particular. In areas of intense folding there
commonly are isolated, tight fold closures sandwiched between apparently unfolded foliation or
layering surfaces; they die out upward and downward in otherwise unfolded rock; such structures
are referred to as intrafolial folds or, if dismembered, as rootless intrafolial folds. Flattened folds
The concept of flattened folds is based on the idea that extreme shortening modifies the shape at
highamplitude buckle folds. Homogeneous flattening thickens the hinge regions, thins the limbs,
and gradually reduces the interlimb angle. Estimates of flattening strain in Class 1C folds allow
restoration of parallel fold shapes that in turn can be used to decipher the amount of shortening
due to buckling. A simple and direct method to estimate flattening strain (the late-stage pure
shear of buckling history) assumes that class1C, 2 or 3 folds initially had a class 1B (concentric
and parallel) shape. The estimate of buckle shortening, in turn, allows restoration of original
length and thickness of undistorted beds, provided the layer-parallel shortening during the initial
stages of folding is insignificant. Several techniques allow determining the amount of flattening,
with the assumption that strain was homogeneous. However, none of these techniques considers
the layer shortening that precedes buckling and therefore each gives a minimum bulk strain
associated with the fold. An additional limitation is that all are two-dimensional. t ′ / α method:
The amount of strain is determined from the way the orthogonal thickness t and the thickness
parallel to the axial plane T of folded layers vary as a function of the limb dip. Inverse thickness
method: The thickness of the layer at any point around the fold is inversely proportional to the
stretch (final length / original length) of the tangent to the folded layer at the angle of dip at
which the thickness is measured. The technique is simple. (1) Measure the orthogonal
thicknesses t perpendicular to tangents drawn to folded layer. (2) Plot inverse thicknesses 1/t
from a common point (i.e. in polar coordinates), each in the direction of the tangent line. A strain
ellipse emerges that discloses strain due to flattening. Folds jpb, 2017 221 Fold-center to
median-layer-line distance: The method is based on the assumption that, in concentric folds, the
center to layer length in any direction is the diameter of an ellipse of the same aspect ratio as the
strain ellipse. (1) On the fold profile, define the center as the intersection between the axial plane
and the median line, which joins inflexion points. (2) Draw lines at regular angular spacing
through the center. (3) Measure the distance d from the center to the middle of the folded layer
along each line. Again, a strain ellipse emerges on a polar graph where d is plotted as a function
of line orientation. Isogon rosette: Dip isogons are drawn on the profile section of a fold by
linking the points of equal dip on the inner and outer arcs. They can be arranged in a rosette by
displacing the isogons without changing their orientation until the mid-point of each isogon
becomes the common point of intersection of all isogons. The end points of isogons in the rosette
trace a characteristic curve that defines the fold geometry. This curve is a circle in parallel folds,
an ellipse in flattened parallel folds, and it reduces to a pair of points in ‘‘similar’’ folds. Since
isogons deform as material lines during flattening, the characteristic curve, namely, the ellipse,
directly represents the strain ellipse in flattened parallel folds. The ‘‘isogon rosette’’ method
allows representation of a given fold by a point on the RS − θ plot, where Rs and θ are the
twodimensional strain ratio and the angle between the maximum principal strain and the fold
axial trace, respectively. Axial plane continuity - harmonic and disharmonic folds Folds in which
axial planes are continuous across successive folded layers that show approximately the same
wavelength and amplitude are harmonic. Typically, similar folds that ideally maintain their shape
throughout a section are harmonic. Folds in which the amplitude, wavelength and style change
along discontinuous axial surfaces from one layer to another are disharmonic. Disharmonic folds
Folds jpb, 2017 222 develop because of differing rheology in the different layers. The
incompetent beds are squeezed and adapt to the form imposed by the competent beds.
Disharmonic folds are particularly common in areas of parallel folds because they die out with
increasing depth along their axial plane, reaching over-tightening in the core. Consequently their
extent is limited to the centre of curvature beyond which shortening should be accommodated by
faulting or another type of folding. They may terminate downward at some surface, called
detachment or décollement, along which they are separated (decoupled) from unfolded layers
below. The Jura Mountains are typically cited for such relationships. Symmetry Folds are
symmetrical if the axial plane bisects the interlimb angle and divides the fold in two identical
halves. In symmetrical folds with a vertical axial plane the hinge line passes through the crest
and trough of antiforms and synforms, respectively. If the axial plane is not a plane of symmetry,
the limbs have unequal lengths and one limb dips more steeply than the other: the folds are
asymmetric. Their leaning direction suggests a relative sense of movement, termed the apparent
vergence. Fore limb - Back limb An asymmetric or overturned antiform has a steep and short
forelimb and a gentler, longer back limb. The forelimb is stratigraphically inverted in overfold
anticlines. This forelimb is reversed, or inverted or overturned, whilst the back limb is normal.
Note that the overturned limb dips steeper, and the normal limb shallower than the axial plane.
Note also that inverted bedding youngs in the opposite direction to which it dips. Folds jpb, 2017
223 Facing and vergence The direction of apparent movement of the upper, long limb with
respect to the shorter limb of an asymmetric fold is called the vergence. In other word, vergence
is simply the sense of asymmetry. The true vergence or facing is the younging direction along the
axial plane in a direction perpendicular to the fold axis. Uniformly verging asymmetric folds are
characteristic of thrust belts. Vergence is useful in working out the regional direction of transport
and help to fix an observation location on large folds. Antiformal synclines and synformal
anticlines are downward facing folds since the stratigraphy is completely inverted in passing
along the axial plane; conversely, anticlines and synclines are upward facing folds. Downward
facing folds are commonly formed during refolding of the stratigraphically inverted limb of a
recumbent fold. Attention: in the German literature the term has been applied to the direction
toward which the fold is leaning, with the idea that the form of the fold profile represents the
frozen movement and rotation that generated this fold form. The movement picture is supposable
where all folds express a unique one-sided regional movement. However, the asymmetry of
subsidiary folds varies from one limb of a large fold to the other limb (see paragraph on parasitic
folds, below). Strictly speaking, the form of a fold should be defined as an apparent vergence by
opposition to the true vergence that includes the younging direction. The true vergence of first
order folds then defines the movement picture. Fold trains A fold train is a series of folds along a
particular layer or series of layers. Several folds may develop next to each other within a soft
layer between virtually unfolded competent layers. When such layer-bounded fold trains display
a systematic vergence, the sense of fold asymmetry affords a bulk relative sense of layer-parallel
shear. Such folds are called drag folds, the implication being that the shear component of the
velocity gradient across the layers has dragged the soft layer into a suite of folds
characteristically non-cylindrical, asymmetric and disharmonic (i.e. the soft layer becomes
detached from the adjacent layers). In the same way, drag folds may also develop within a thrust
zone. Gravitational forces acting on plastically deforming layers can produce cascades of folds.
Parasitic folds Hinge zones and limbs of large folds often display folds of smaller wavelength
and amplitude: larger and smaller folds are together polyharmonic. The small folds are called
parasitic or subsidiary folds with respect to the larger ones. The largest folds are termed
first-order folds, the next largest are called second-order folds and so forth. The axes of parasitic
folds are habitually closely parallel to the axis of the major fold with which they are associated.
They are said to be congruous, by contrast with incongruous parasitic folds whose axes deviate
appreciably from the attitude of the major fold axis. Folds jpb, 2017 224 First-order folds may be
symmetrical whereas the congruous, second-order folds are asymmetrical. The sense of
asymmetry, referred to as local apparent vergence, is consistently towards the hinges of
higher-order antiforms. It changes systematically across the axial surfaces of the first-order folds,
so that, looking down the axial direction, all parasitic folds in a limb have a clockwise apparent
vergence and are described as Z folds, whereas those in the other limb have anticlockwise
apparent vergence and are S folds. Symmetrical M folds generally occur in the hinge zone (W
may be used for synforms opposed to antiforms). The axial plane of the first order fold links and
runs across second order M folds and separates limbs with S and Z second-order folds. In the
field, the asymmetry is used to locate hinges of the next larger order folds if they were both
generated together. Note that flexural slip and/or flow provides a coherent explanation of why Z
and S second order folds form on the limbs of a first order fold. Parasitic folds would initiate as
symmetric buckle folds sheared hingewards during flexural flow of incompetent layers.
Enveloping and median surfaces The average orientation of a folded surface affected by a fold
train is measured from the enveloping surface, which is constructed either as tangential to most
or all of the hinge zones in the folded surface. The enveloping surface defines the limits of folds,
thus relates the geometry of small- to large-scale folds in areas where there are so many
small-scale folds that they obscure the general orientation of bedding (e.g. in a tightly folded
limb). Which folds are touched by the enveloping surface depends on the scale at which the
structure is being considered. The median surface joins the successive inflection lines of a folded
surface. The median surface separates antiforms from synforms. The median and enveloping
surfaces are generally almost parallel and, therefore, yield the same information. Axial planes of
symmetrical folds are perpendicular to the enveloping and median surfaces. Wavelength, arc
length, amplitude and aspect ratio Amplitude and wavelength define the size of a single fold and
refer to the mathematical terminology used to describe a sinusoidal curve. Wavelength The
distance between two successive anticlinal (or synclinal) hinges seen in profile is the wavelength.
The distance between inflexion points can also be used to measure the wavelength. The arc
length is the distance along the folded plane between two points separated by one wavelength.
Amplitude The amplitude is measured by taking half the distance along the axial plane from one
anticlinal hinge to the surface enveloping the two adjoining synclinal hinges (or vice versa), in
other words, the distance along the axial plane from the median surface to the hinge. A train of
folds may fold a surface periodically or non-periodically. The term pericline is applied to
large-scale antiforms or synforms whose amplitude decreases regularly to zero in both directions,
so that the fold has precise limits in space. Domes and basins are periclinal structures. Aspect
ratio The ratio of the amplitude to half the wavelength is the aspect ratio of a fold. Where folding
is disharmonic, both wavelength and amplitude of folds vary between successive layers.
Conjugate folds The term conjugate is used to describe any pair of identical folds that have axial
surfaces inclined at a high angle to one another in profile (resulting in opposite vergence) or in
map view. Such folds generally terminate in an angular fold, of yet another orientation, at the
point where their axial surfaces meet or intersect one another, in general in an incompetent layer.
Conjugate folds with round hinge zones are box folds. Conjugate kink bands (conjugate folds
with angular hinges) are common. Collapse folds In regions folded in near surface conditions,
folds with rather flat-lying axial planes and a local vergence opposite to that of parasitic folds are
due to gravity-driven collapse into the synclinal regions of layers first steepened by folding.
Flaps are overturned sequences that result from bent over backward gravitational instabilities,
without breaking. Morphological classification - Thickness variation of a folded layer Structural
geologists have long used the shapes of folds for determining the amount of shortening due to
folding in layered rocks. They have elaborated a range of techniques to estimate shortening from
fold shapes. Geometrical parameters The variable curvatures of hinge zones and limbs of
successive surfaces reflect thickness variations of layers resulting from folding. These shape
changes can be categorised in terms of three interdependent measurements on the profile: α = dip
angle at different points on successive folded surfaces; tα = layer thickness normal to layering
where α is measured; Tα = layer thickness parallel to the axial plane where α is measured. These
three measurements are the basis of a classification that refers to dip isogon patterns: A dip
isogon is a line linking points of equal limb dip on the outer and inner arcs of a folded layer seen
in profile section (i.e. seen orthogonal to the fold axis). This classification involves an indirect
relationship everywhere on the fold: Tα∗cosα = tα The thickness at the fold hinge is t 0 = T0 .
The ratio t ′ = tα / t0 or T’=Tα / T0 can be calculated and plotted versus the angle α. This ratio
expresses changes in orthogonal thickness with change in dip α . Construction Dip isogons are
constructed as follows: - Draw the trace of the axial plane on a profile view of the fold (i.e.
orthogonal to the fold axis). - Draw another line perpendicular to this trace, preferably out of the
fold. - With a protractor gliding along the line orthogonal to the axial plane, locate points along
the folded surface whose tangents intersect the glide-line at specific angles. - A set of iso-angle
points on adjacent folded surfaces permits the drawing of dip isogons. The orientations of the dip
isogons over a fold qualitatively describe the variation in thickness and the difference of
curvatures between successive interfaces. If isogons are convergent towards the core of the fold,
the curvature of the outer arc is less than that of the inner arc. Conversely, if isogons are
divergent towards the core of the fold, the curvature of the outer arc is greater than that of the
inner arc. Fold types The characteristics of the dip isogons leads to a three-fold classification,
with two additional subclasses: Class 1 folds with isogons convergent towards the core of the
fold (curvature of the inner arc is greater than outer arc), class 2 folds with isogons parallel to the
axial trace (curvature of the inner arc is equal to outer arc) and class 3 folds with divergent
isogons (curvature of the inner arc is smaller than outer arc). Class 1 is subdivided into three
sub-classes according to the degree of convergence: - class 1A are strongly convergent, meaning
that thickness of folded layers in hinges are thinner than those in limbs. - class 1B are a special
case of parallel folds where the isogons are perpendicular to the fold surface. - class 1C folds are
weakly convergent, meaning that thicknesses of folded layers in hinges are thicker than those in
limbs. Class 2 folds are a special case of similar folds along which the curvature of successive
folded surfaces remains the same. Class 3 are folds where the curvature of successive folded
surfaces decreases toward the inner arc of the fold. These classes are shown on the t ′ / α plot.
This classification is purely geometric and tells nothing about folding processes. Folds jpb, 2017
228 Fault-related folds There commonly are folds geometrically associated with faults. They are
in general controlled by the fault geometry. Drag folds Bedding is frequently bent in fault zones.
These local flexures are known as drag folds and fault drags. They usually are convex toward the
direction of the fault movement and are thus attributed to some frictional resistance to slip along
the fault plane. This interpretation suggests that faulting is initiated first and that folding occurs
adjacent to the fault as one block is dragged along the other (normal drag). However, folding
might precede faulting, drag folds representing bending of rock before it breaks. Folds jpb, 2017
229 The use of drag folds, intuitively inferring that the direction of folding is toward the
direction of fault movement, can be misleading because convexity opposite to the sense of
displacement, termed reverse drag, is common. For example, roll-over flexures on listric normal
faults are hanging wall folds concave towards the slip direction. Reverse drag is hardly
distinguished from normal drag when they appear separately. In addition, the orientation of drag
folds is often not controlled by the movement direction but rather the intersection between
bedding and the fault plane, and the drag may vary from reverse to normal from the centre to the
termination of faults. Therefore, drags should be used with extreme care to ascertain the sense of
slip along faults. Trains of drag folds are common in incompetent layers between two competent
layers of in the proximity of thrust faults. En échelon folds In some non-cylindrically folded
surfaces, doubly plunging and relatively short, nearly upright folds in parallel series have
alternating antiform and synform axes oblique to the fold string. Such folds are stepped and
consistently overlapping; they define an en échelon array. Note, however, that this term describes
the geometry of the folded surface and is independent of the relationship of the structure to the
horizontal and vertical. Taking the steep axial planes as roughly orthogonal to the shortening
direction, their distribution permits to decipher the potential fault they are related to. Such folds
are common above strike-slip faults that have not broken the cover but offset the basement
blocks. The en échelon fold-pattern reveals the relative sense of movement along the basement
fault. Drape folds - Forced folds Cover rocks can be more or less passively bent to conform to
the topography of the buried basement / cover interface. An important shape-controlling factor is
whether the cover remains welded to or is detached from the basement. Drape folds are generally
open curvatures in a sedimentary layer that conforms passively to the configuration of
underlying structures and geological bodies. A fold formed by differential compaction is an
example. Forced folds are generally fault-related, long and linear flexures that relative
movements of basement blocks generate in cover rocks. Their overall shape and trend are
dominated by the shape and trend of the underlying forcing fault blocks. They are typically
monoclines with long, gently-dipping backlimbs and short, steeply-dipping forelimbs, the latter
overlying the fault surface. They are equally common in compression and extension regimes.
The type and amount of fault movement controls the fold profile geometry. Thrust-related folds
Thrust movement at depth generates geometrically necessary folds in the allochthonous
hanging-wall as it moves over topographic irregularities of the thrust faults. Kink-like, box folds
result. Two types of ramp-related folds are common in thrusts belts. - Fault bend folds form and
grow above a footwall ramp where a fault steps from a lower flat to a higher one and strata of the
hanging wall slide over the fault bend. As slip occurs, the hangingwall is folded to accommodate
to the shape of the footwall ramps in a passive syncline-anticline pair at the base and top of the
ramp. Specific geometries are maintained throughout the development of the anticline. (i) The
ramp anticline terminates downward into the upper-flat. (ii) The backlimb is parallel to the
footwall ramp. (iii) The forelimb is shorter and steeper than the backlimb. (iv) the
anticlinesyncline pair directly reflects the geometry of the fault bends. - In a fault-propagation
fold the ramp does not continue to an upper flat. Strata cut by the base of the ramp are shortened
by thrusting. Fault slip decreases to zero in the up-section direction and the fault dies out into the
axial surface of a syncline. Strata above and in advance of the upper tip line of the propagating
thrust are shortened entirely by folding. Typically (i) such folds are asymmetric in the direction
of thrust movement, (ii) they tighten with increasing displacement, and (iii) both limbs lengthen
while the fault tip propagates upwards. - Growth folds develop in sedimentary strata at the same
time as they are being deposited. Antiformal structures commonly grow over a ramp or a duplex
zone to build antiformal stacks. The common association of folds and thrusts at a regional scale
defines a fold and thrust belt. The fold shape is determined by the shape of an advancing ramp
that does not tie into an upper flat. The ramp fault is replaced upward by an asymmetric fold,
which is overturned in the direction of transport. There is a systematic and predictable geometric
relation between a fold and the thrust that generated it, and we can therefore use fold geometry to
infer fault position and geometry at depth. Normal-fault-related folds As in thrust systems,
geometrically necessary folds are generated above topographic irregularities of extensional
faults. Tilting of the hanging wall of a normal listric fault toward the main fault produces a
half-antiform called rollover anticline. The rollover anticline is a gentle convex bending of
upperblock beds necessary to accommodate the concavity of a listric normal fault. As the
hanging-wall slips along the deep, gently dipping parts of the fault, an additional volume is
generated between the footwall and the hanging wall along the shallower, steeper parts of the
fault. Beds initially horizontal in the hanging wall must become gently convex upwards to fill up
this additional space. Tear faults In areas like the Jura, strike-slip faults seem to tear the folds
across. They are attributed to differential advance of adjoining segments of the folds, with folds
in one block being more closed and tighter than in the other. Tear- faults as such are transfer
faults that originated during folding. Fracture patterns in folds Geometry and density of fractures
accommodate the strain induced in the strata during folding. Fracture sets are commonly
documented to vary from hinge to limb. - In hinge regions fractures are mostly parallel to the
fold axis and orthogonal to bedding. Joints and normal faults are formed by curvature-related,
extensional stresses within the outer-arc of the fold. Stylolitic joints and thrusts are formed by
compressional stresses within the inner-arc of the fold. - In limbs fractures are mostly parallel to
the fold axis but oblique to bedding. They are related to interbed slip. Primary folds Lavas
commonly develop surface folds during flow. These folds are due to the different properties of
the chilled outer crust of the lava relative to the hot, faster flowing inner part. In magmatic rocks,
folds often deform the primary banding during the late-stage flow of the magma. Non-lithified
sediments are considered to be suspensions that can deform in a ductile manner. Gravity-driven
sliding (slumping) of unconsolidated layers therefore often produces complex fold shapes called
in short slumps. Slumped layers typically lay between undeformed strata. Determining a shear
direction from fold orientations Layers in sheared rocks commonly form non-cylindrical
drag-folds whose axial planes are near the shear plane, sheath folds being one end member of
this geometry. The hinges are curved because their orientation depends on several parameters
such as the original orientation of the layer relative to the shear plane and local heterogeneities of
flow. Furthermore, fold hinges can form parallel to the shear direction, or can rotate towards the
shear direction during progressive deformation. Therefore, fold hinges are rarely orthogonal to
the slip direction. Because of such important local changes, isolated small scale folds alone
frequently give misleading transport directions. A safe assumption is that the asymmetry of any
fold pertaining to the same set is consistent with the sense of shear that produced them. This
consistency is the basis of the geometric method that provides a guide to the slip direction. The
(Hansen) method considers a group of minor folds and proceeds as follows: 1) All hinge
orientations are plotted on a stereonet with their individual sense of asymmetry. 2) All hinges
should lie approximately on the same axial plane (e.g. a great circle), near the shear plane. 3) The
separation arc of the great circle across which the sense of asymmetry becomes opposite contains
the slip direction, approximately along the bisector of the angle defining this arc. 4) The
asymmetry of folds defines the bulk sense of shear. Conclusion A fold is a bend in a layered rock
caused by compressive stress (buckling) or passive draping of layers over a lower structure or
around a resistant object. Folds represent large-scale flow of material. Folds display a wide range
of shapes and result from a wide range of processes that both largely reflect the rock behaviour.
It is therefore common to observe that geometrical characteristics change within the same fold
from layer to layer. Antiforms are hydrocarbon reservoirs. Folds host ore deposits in hinge areas
due to flow of material to those localities. Folds form associated with faults and thus can signal
earthquake hazards. Folds record periods of rock deformation, therefore generations need to be
distinguished and dated. For these reasons, structural geologists have to document the fold
geometry for economic and geologic hazard applications, the layer configuration as clues to
conditions of deformation and the definition of fold generations for geologic history.
Intrusion Types
Intrusions are also classified according to size, shape, depth of formation, and geometrical
relationship to the country rock. Intrusions that formed at depths of less than 2 kilometers are
considered to be shallow intrusions, which tend to be smaller and finer grained than deeper
Dikes. A dike is an intrusive rock that generally occupies a discordant, or cross‐cutting, crack or
fracture that crosses the trend of layering in the country rock. Dikes are called pegmatites when
they contain very coarse‐grained crystals—a single such crystal can range in size from a few
centimeters to 10 meters in diameter.
Sills. Sills are formed from magmas that entered the country rock parallel to the bedding
(layering) and are thus concordant with the country rock. Sills can sometimes look like volcanic
flows that were interbedded with sedimentary units.
Laccoliths. A laccolith resembles a sill but formed between sedimentary layers from a more
viscous magma that created a lens shaped mass that arched the overlying strata upward.
Volcanic necks. A volcanic neck is the rock that formed in the vent of a volcano at the end of its
eruptive life and remains “standing” after the flanks of the volcano have eroded away.
Plutons. Plutons are discordant intrusive
rocks that formed at great depths. They tend
to be large, coarse grained, and irregular in
shape. If the intrusion occupies less than
100 square kilometers (60 square miles) at
the earth's surface it is called a stock; if it is
larger than 100 square kilometers, it is
termed a batholith. Batholiths are usually
composed of granite. They have formed
over long periods through the accumulation
of smaller magma blobs called diapirs,
which result from localized melting of the
crust; the diapirs then slowly move upward
toward the surface and coalesce into a larger
mass. Granitic batholiths usually form the
cores of mountain complexes and are a result of plate tectonic action.
In most cases, a body of hot magma is less dense than the rock surrounding it, so it has a
tendency to creep upward toward the surface. It does so in a few different ways:
● Filling and widening existing cracks
● Melting the surrounding rock (called country rock)
● Pushing the rock aside (where the rock is hot enough and under enough pressure to
deform without breaking)
● Breaking the rock.
When magma forces itself into cracks, breaks off pieces of rock, and then envelops them, this is
called stoping. The resulting fragments are called xenoliths.
Plutons can have different shapes and different relationships with the surrounding country rock.
These characteristics determine what name the pluton is given.
Large, irregularly shaped plutons are called stocks or batholiths, depending on size. Tabular
plutons are called dikes if they cut across existing structures, and sills if they do not. Laccoliths
are like sills, except they have caused the overlying rocks to bulge upward. Pipes are cylindrical
Intrusions can be classified according to the shape and size of the intrusive body and its relation
to the other formations into which it intrudes:
Batholith: a large irregular discordant intrusion
Chonolith: an irregularly-shaped intrusion with a demonstrable base
Cupola: a dome-shaped projection from the top of a large subterranean intrusion
Dike: a relatively narrow tabular discordant body, often nearly vertical
Laccolith: concordant body with roughly flat base and convex top, usually with a feeder pipe
Lopolith: concordant body with roughly flat top and a shallow convex base, may have a feeder
dike or pipe below
Phacolith: a concordant lens-shaped pluton that typically occupies the crest of an anticline or
trough of a syncline
Volcanic pipe or volcanic neck: tubular roughly vertical body that may have been a feeder vent
for a volcano
Sill: a relatively thin tabular concordant body intruded along bedding planes
Stock: a smaller irregular discordant intrusive
Boss: a small stock
Deposition is the process that follows erosion. Erosion is the removal of particles (rock, sediment
etc.) from a landscape, usually due to rain or wind. Deposition begins when erosion stops; the
moving particles fall out of the water or wind and settle on a new surface. This is deposition.
The overall cause for deposition is erosion, since the particles need to be moving in order to stop.
However, there has to be something that causes the erosion to stop and the deposition to begin.
This transition is caused by a change in the agent of transport. Water can slow or evaporate,
allowing sediment to stop being carried along. Wind can die down and release soil. Ice can melt
and release its hold. Any such change begins the process of deposition.
Erosion can be a very destructive force, but together with deposition, it can also be a force of
creation. These two processes are responsible for the creation of new landscapes, including hills,
valleys and coastlines. Though erosion can alter an area, the affected parts are not destroyed but
simply moved. Deposition allows these parts to settle elsewhere.
Deposition is the addition of sediment to a landform. Landforms are weathered and eroded all the
time. Of course, their pieces have to go somewhere. They’ll be carried for a while by wind,
water, or some other force. Yet, eventually, the sediment will be deposited onto a different
Rocks that have been depoited at the
base of a slope. The resulting
landform is known as a scree. Generally, screes form as a result of deposition by mass wasting.
The Science of Deposition
Deposition seems like a pretty simple concept, right? Things get picked up and eventually they
get put down. That’s basically just how gravity works!
Well, things are actually a bit more complicated than that. A variety of forces move sediment in
both patterned and messy ways. This can produce clear landforms, like deltas. Yet, it can also
give us areas of plain sediment, like sandy deserts. The existence of these unique land features
tells us that deposition isn’t totally random.
But first, we have to know where deposited sediments come from. In this article, we’re mainly
going to focus on the ‘putting down’ side of sediment movement (deposition). You can read
more about how sediment is picked up (weathering/erosion) in the articles to the left. For now,
all we need to know is that weathering and erosion are part of the rock cycle. Their role in it is to
break large rocks into the sediment. Outside of the rock cycle, they also take materials from soil
Surface runoff flowing
over some soil. It breaks
up the soil into smaller
(sediment) have to settle
Whenever they do, we call
it deposition. Deposition
creates both interesting
and important landforms.
A few of them, like screes,
mysterious. Others are
central to how humans
beings live! With that, let’s explore a few examples of deposition.
Examples and Pictures of Deposition
Soil Deposition
A central concern for farmers is finding fertile soil for plant growth. They’re well aware that soil
comes in many different types. That’s because it’s created by erosion and deposition. Soil is
actually made-up of a mixture of different sediments, like weathered rock and clay.
Getting a good mixture for farming (and in the right order) can actually take thousands of years!
For this reason, it’s really important that we can find good soil. Fortunately, the science of
deposition tells us where to locate it.
Plants growing in a patch of
soil. You might think that’s kind
of boring. But, good soil is
hard to come by!
One of the best kinds of soil
for farming is silty soil. Silt is a
very fine sediment that’s made
when mineral-rich rocks are
weathered. It’s often produced
by wind erosion which is pretty
environments. Then, during
floods, the silt is picked up and
deposited near topographical
low points. These points are
usually rivers or coastlines.
The Nile River. Even
though it’s in a desert, it’s
a great place for farming.
As a result, the land
around that water is some
of the best for farming.
This is true even in
deserts! But, we would
only know that if we
understood how silt is deposited. In a sense, we need the science of deposition to produce food.
And what else is more important than that?
Deposition of Rocks
If you’ve ever gone hiking, you might have seen lots of little rock piles laying around. Who put
those there? A Sasquatch?
Obviously not. Instead, these piles were deposited by natural forces. As we said in the beginning,
mass wasting tells us that rocks gradually fall or slide downhill. When they hit the ground, they
might break into small pieces or piles.
Rocks are deposited in a number of other ways, too. For example, eroded rocks often find their
way into rivers. If they’re small enough, the river will carry them downstream. All the way at the
end, they may be deposited in a delta. A delta is a landform between a flowing body of water
(here, a river) and one that’s not moving. They tend to contain a ton of sediment.
A river delta. You can
see brownish sediment
throughout the water.
In earth science, deformation is an alteration of the size or shape of rocks. Deformation is caused
by stress, the scientific term for force applied to a certain area. Stresses on rocks can stem from
various sources, such as changes in temperature or moisture, shifts in the Earth’s plates, sediment
buildup or even gravity.
Types of Deformation
There are three types of rock deformation. Elastic deformation is temporary and is reversed when
the source of stress is removed. Ductile deformation is irreversible, resulting in a permanent
change to the shape or size of the rock that persists even when the stress stops. A fracture or
rupture, also known as brittle deformation, results in the breakage of the rock. Like ductile
deformation, fractures are irreversible.
Factors and Examples
Certain factors determine which type of deformation rock will exhibit when exposed to stress.
These factors are rock type, strain rate, pressure and temperature. For instance, higher
temperatures and pressures encourage ductile deformation. This is common deep within the
Earth, where, due to higher temperatures and pressure than nearer the surface, rocks tend to be
more ductile.
Stress and Strain
We start our discussion with a brief review of the concepts of stress and strain. Recall that stress
is a force acting on a material that produces a strain. Stress is a force applied over an area and
therefore has units of Force/area (like lb/in 2 ). Pressure is a stress where the forces act equally
from all directions.
If stress is not equal from all directions then we say that the stress is a differential stress. Three
kinds of differential stress occur.
1. Tensional stress (or extensional stress), which stretches rock;
2. Compressional stress, which squeezes rock; and
3. Shear stress, which result in slippage and translation.
When rocks deform they are said to strain. A strain is a change in size, shape, or
volume of a material. We here modify that definition somewhat to say that a strain
also includes any kind of movement of the material, including translation and tilting.
Stages of Deformation
When a rock is subjected to increasing stress it passes through 3 successive stages of
Elastic Deformation -- wherein the strain
is reversible
Ductile Deformation -- wherein the strain
is irreversible.
Fracture - irreversible strain wherein the
material breaks.
We can divide materials into two classes that depend on their relative behavior under
Brittle materials have a small or large region of elastic behavior but only a small
region of ductile behavior before they fracture.
● Ductile materials have a small region of elastic behavior and a large region of ductile
behavior before they fracture.
How a material behaves will depend on several factors. Among them are:
Temperature - At high temperature molecules and their bonds can stretch and move,
thus materials will behave in more ductile manner. At low Temperature, materials
are brittle.
● Confining Pressure - At high confining pressure materials are less likely to fracture
because the pressure of the surroundings tends to hinder the formation of fractures.
At low confining stress, material will be brittle and tend to fracture sooner.
● Strain rate -- At high strain rates material tends to fracture. At low strain rates more
time is available for individual atoms to move and therefore ductile behavior is
● Composition -- Some minerals, like quartz, olivine, and feldspars are very brittle.
Others, like clay minerals, micas, and calcite are more ductile This is due to the
chemical bond types that hold them together. Thus, the mineralogical composition
of the rock will be a factor in determining the deformational behavior of the rock.
Another aspect is presence or absence of water. Water appears to weaken the
chemical bonds and forms films around mineral grains along which slippage can
take place. Thus wet rock tends to behave in ductile manner, while dry rocks tend to
behave in brittle manner.
Brittle-Ductile Properties of the Lithosphere
We all know that rocks near the surface of the Earth behave in a brittle manner. Crustal
rocks are composed of minerals like quartz and feldspar which have high strength,
particularly at low pressure and temperature. As we go deeper in the Earth the strength of
these rocks initially increases.
At a depth of about 15 km we reach a point
called the brittle-ductile transition zone.
Below this point rock strength decreases
because fractures become closed and the
temperature is higher, making the rocks
behave in a ductile manner. At the base of the
crust the rock type changes to peridotite which
is rich in olivine. Olivine is stronger than the
minerals that make up most crustal rocks, so
the upper part of the mantle is again strong.
But, just as in the crust, increasing
temperature eventually predominates and at a
depth of about 40 km the brittle-ductile
transition zone in the mantle occurs. Below
this point rocks behave in an increasingly
ductile manner.
Deformation in Progress
Only in a few cases does deformation of rocks occur at a rate that is observable on human
time scales. Abrupt deformation along faults, usually associated with earthquakes occurs
on a time scale of minutes or seconds. Gradual deformation along faults or in areas of
uplift or subsidence can be measured over periods of months to years with sensitive
measuring instruments.
Evidence of Past Deformation
Evidence of deformation that has occurred in the past is very evident in crustal rocks. For
example, sedimentary strata and lava flows generally follow the law of original
horizontality. Thus, when we see such strata inclined instead of horizontal, evidence of an
episode of deformation.
Since many geologic features are planar in nature, we a way to uniquely define the
orientation of a planar feature we first need to define two terms - strike and dip.
Fracture of Brittle Rocks
As we have discussed previously, brittle rocks tend to fracture when placed under a high
enough stress. Such fracturing, while it does produce irregular cracks in the rock,
sometimes produces planar features that provide evidence of the stresses acting at the time
of formation of the cracks. Two major types of more or less planar fractures can occur:
joints and faults.
As we learned in our discussion of physical weathering, joints are fractures in rock that
show no slippage or offset along the fracture. Joints are usually planar features, so their
orientation can be described as a strike and dip. They form from as a result of extensional
stress acting on brittle rock. Such stresses can be induced by cooling of rock (volume
decreases as temperature decreases) or by relief of pressure as rock is eroded above thus
removing weight.
Joints provide pathways for water and thus pathways for chemical weathering attack on
rocks. If new minerals are precipitated from water flowing in the joints, this will form a
vein. Many veins observed in rock are mostly either quartz or calcite, but can contain rare
minerals like gold and silver. These aspects will be discussed in more detail when we talk
about valuable minerals from the earth in a couple of weeks.
Because joints provide access of water to rock, rates of weathering and/or erosion are
usually higher along joints and this can lead to differential erosion.
From an engineering point of view, joints are important structures to understand. Since
they are zones of weakness, their presence is critical when building anything from dams to
highways. For dams, the water could leak out through the joints leading to dam failure.
For highways the joints may separate and cause rock falls and landslides.
Faults occur when brittle rocks fracture and there is an offset along the fracture. When the
offset is small, the displacement can be easily measured, but sometimes the displacement
is so large that it is difficult to measure.
Types of Faults
As we found out in our discussion of earthquakes, faults can be divided into several
different types depending on the direction of relative displacement. Since faults are planar
features, the concept of strike and dip also applies, and thus the strike and dip of a fault
plane can be measured. One division of faults is between dip-slip faults, where the
displacement is measured along the dip direction of the fault, and strike-slip faults where
the displacement is horizontal, parallel to the strike of the fault. Recall the following types
of faults:
Dip Slip Faults - Dip slip faults are faults that have an inclined fault plane and
along which the relative displacement or offset has occurred along the dip
direction. Note that in looking at the displacement on any fault we don't know
which side actually moved or if both sides moved, all we can determine is the
relative sense of motion.
Normal Faults - are faults that result from horizontal tensional stresses in
brittle rocks and where the hanging-wall block has moved down relative to
the footwall block.
Horsts & Grabens - Due to the tensional stress responsible for normal faults, they often
occur in a series, with adjacent faults dipping in opposite directions. In such a case the
down-dropped blocks form grabens and the uplifted blocks form horsts. In areas where
tensional stress has recently affected the crust, the grabens may form rift valleys and the
uplifted horst blocks may form linear mountain ranges. The East African Rift Valley is an
example of an area where continental extension has created such a rift. The basin and
range province of the western U.S. (Nevada, Utah, and Idaho) is also an area that has
recently undergone crustal extension. In the basin and range, the basins are elongated
grabens that now form valleys, and the ranges are uplifted horst blocks.
Reverse Faults - are faults that result from horizontal compressional stresses in
brittle rocks, where the hanging-wall block has moved up relative the footwall
A Thrust Fault is a special case of a reverse fault where the dip of the fault is less than
45o. Thrust faults can have considerable displacement, measuring hundreds of kilometers,
and can result in older strata overlying younger strata.
Strike Slip Faults - are faults where the relative motion on the fault has taken
place along a horizontal direction. Such faults result from shear stresses acting in
the crust. Strike slip faults can be of two varieties, depending on the sense of
displacement. To an observer standing on one side of the fault and looking across
the fault, if the block on the other side has moved to the left, we say that the fault is
a left-lateral strike-slip fault. If the block on the other side has moved to the right,
we say that the fault is a right-lateral strike-slip fault. The famous San Andreas
Fault in California is an example of a right-lateral strike-slip fault. Displacements
on the San Andreas fault are estimated at over 600 km.
Synclines are folds where the originally
horizontal strata have been folded
downward, and the two limbs of the
fold dip inward toward the hinge of the
fold. Synclines and anticlines usually
occur together such that the limb of a
syncline is also the limb of an anticline.
In the diagrams above, the fold axes are horizontal, but if the fold axis is not
horizontal the fold is called a plunging fold and the angle that the fold axis makes
with a horizontal line is called the plunge of the fold.
Note that if a plunging fold intersects
a horizontal surface, we will see the
pattern of the fold on the surface (see
also figures 11.15e in your text.
Domes and Basins are formed as a result of vertical crustal motion. Domes look like an
overturned bowl and result from crustal upwarping. Basins look like a bowl and result
from subsidence (see figure 11.14 in your text).
Folds are described by the severity of folding. an open fold has a large angle between
limbs, a tight fold has a small angle between limbs.
Further classification of folds include:
If the two limbs of the fold dip away from the axis with the same angle, the fold is
said to be a symmetrical fold.
If the limbs dip at different angles, the folds are said to be asymmetrical folds.
If the compressional stresses that cause the folding are intense, the fold can close
up and have limbs that are parallel to each other. Such a fold is called an isoclinal
fold (iso means same, and cline means angle, so isoclinal means the limbs have the
same angle). Note the isoclinal fold depicted in the diagram below is also a
symmetrical fold.
If the folding is so intense that the strata on one limb of the fold becomes nearly
upside down, the fold is called an overturned fold.
An overturned fold with an axial plane that is nearly horizontal is called a
recumbant fold.
A fold that has no curvature in its hinge and straight-sided limbs that form a zigzag
pattern is called a chevron fold.
Folds and Topography
Since different rocks have different resistance to erosion and weathering, erosion of folded
areas can lead to a topography that reflects the folding. Resistant strata would form ridges
that have the same form as the folds, while less resistant strata will form valleys (see
figure11.14 in you text).
How Folds Form
Folds develop in two ways:
Flexural folds form when layers slip as stratified rocks are bent. This results in the
layers maintaining their thickness as they bend and slide over one another. These
are generally formed due to compressional stresses acting from either side.
Flow folds form when rocks are very ductile and flow like a fluid. Different parts
of the fold are drawn out by this flow to different extents resulting in layers
becoming thinner in some places and thicker in outer places. The flow results in
shear stresses that smear out the layers.
Folds can also form in
relationship to faulting of
other parts of the rock body.
In this case the more ductile
rocks bend to conform to the
movement on the fault.
Also since even ductile rocks can
eventually fracture under high
stress, rocks may fold up to a certain
point then fracture to form a fault.
Folds and Metamorphic Foliation
As we saw in our discussion of metamorphic rocks, foliation is a planar fabric that
develops in rocks subject to compressional stress during metamorphism. It may be
present as flattened or elongated grains, with the flattening occurring perpendicular to the
direction of compressional stress. It also results from the reorientation, recrystallization, or
growth of sheet silicate minerals so that their sheets become oriented perpendicular to the
compressional stress direction. Thus, we commonly see a foliation that is parallel to the
axial plane of the fold.
Shearing of rock during metamorphism can also draw out grains in the direction of shear.
Mountains and Mountain Building Processes
One of the most spectacular results of deformation acting within the crust of the Earth is
the formation of mountain ranges. Mountains frequently occur in elongate, linear belts.
They are constructed by tectonic plate interactions in a process called orogenesis.
Mountain building (orogenesis) involves
Structural deformation.
Igneous Processes.
Constructive processes, like deformation, folding, faulting, igneous processes and
sedimentation build mountains up; destructive processes like erosion and glaciation, tear
them back down again.
Mountains are born and have a finite life span. Young mountains are high, steep, and
growing upward. Middle-aged mountains are cut by erosion. Old mountains are deeply
eroded and often buried. Ancient orogenic belts are found in continental interiors, now far
away from plate boundaries, but provide information on ancient tectonic processes. Since
orogenic continental crust generally has a low density and thus is too buoyant to subduct,
if it escapes erosion it is usually preserved.
Uplift and Isostasy
The fact that marine limestones occur at the top of Mt. Everest, indicates that deformation
can cause considerable vertical movement of the crust. Such vertical movement of the
crust is called uplift. Uplift is caused by deformation which also involves thickening of
the low density crust and, because the crust "floats" on the higher density mantle, involves
another process that controls the height of mountains.
The discovery of this process and its consequences involved measurements of gravity.
Gravity is measured with a device known as a gravimeter. A gravimeter can measure
differences in the pull of gravity to as little as 1 part in 100 million. Measurements of
gravity can detect areas where there is a deficiency or excess of mass beneath the surface
of the Earth. These deficiencies or excesses of mass are called gravity anomalies.
A positive gravity anomaly indicates that an excess of mass exits beneath the area. A
negative gravity anomaly indicates that there is less mass beneath an area.
Negative anomalies exist beneath mountain ranges, and mirror the topography and crustal
thickness as determined by seismic studies. Thus, the low density continents appear to be
floating on higher density mantle.
The protrusions of the crust into the mantle are referred to as crustal roots. Normal crustal
thickness, measured from the surface to the Moho is 35 to 40 km. But under mountain
belts crustal thicknesses of 50 to 70 km are common. In general, the higher the
mountains, the thicker the crust.
What causes this is the principal of isostasy. The principal can be demonstrated by floating
various sizes of low density wood blocks in your bathtub or sink. The larger blocks will
both float higher and extend to deeper levels in the water and mimic the how the
continents float on the mantle (see figure 11.26 in your text).
It must be kept in mind, however that it's not just the crust that floats, it's the entire
lithosphere. So, the lithospheric mantle beneath continents also extends to deeper levels
and is thicker under mountain ranges than normal. Because the lithosphere is floating in
the asthenosphere which is more ductile than the brittle lithosphere, the soft asthenosphere
can flow to compensate for any change in thickness of the crust caused by erosion or
The Principle of isostasy states that there is a flotational balance between low density
rocks and high density rocks. i.e. low density crustal rocks float on higher density mantle
rocks. The height at which the low density rocks float is dependent on the thickness of the
low density rocks. Continents stand high because they are composed of low density rocks
(granitic composition). Ocean basins stand low, because they are composed of higher
density basaltic and gabbroic rocks.
Isostasy is best illustrated by effects of glaciation. During an ice age crustal rocks that are
covered with ice are depressed by the weight of the overlying ice. When the ice melts, the
areas previously covered with ice undergo uplift.
Mountains only grow so long as there are forces causing the uplift. As mountains rise,
they are eroded. Initially the erosion will cause the mountains to rise higher as a result of
isostatic compensation. But, eventually, the weight of the mountain starts to depress the
lower crust and sub-continental lithosphere to levels where they start to heat up and
become more ductile. This hotter lithosphere will then begin to flow outward away from
the excess weight and the above will start to collapse.
The hotter rocks could eventually partially melt, resulting in igneous intrusions as the
magmas move to higher levels, or the entire hotter lower crust could begin to rise as a
result of their lower density. These processes combined with erosion on the surface result
in exhumation, which causes rocks from the deep crust to eventually become exposed at
the surface.
Causes of Mountain Building
There are three primary causes of mountain building.
1. Convergence at convergent plate boundaries.
2. Continental Collisions.
3. Rifting
● Convergent Plate Margins
When oceanic lithosphere subducts beneath continental lithosphere magmas
generated above the subduction zone rise, intrude, and erupt to form volcanic
mountains. The compressional stresses generated between the trench and the
volcanic arc create fold-thrust mountain belts, and similar compression behind
the arc create a fold-thrust belt resulting in mountains. Mountains along the
margins of western North and South America, like the Andes and the Cascade
range formed in this fashion.
Island arcs off the coast of continents can get pushed against the continent.
Because of their low density, they don't subduct, but instead get accreted to the
edge of the continent. Mountain ranges along the west coast of North America
formed in this fashion (see figure 11.20 in your text).
● Continental Collisions
Plate tectonics can cause continental crustal blocks to collide. When this occurs
the rocks between the two continental blocks become folded and faulted under
compressional stresses and are pushed upward to form fold-thrust mountains.
The Himalayan Mountains (currently the highest on Earth) are mountains of this
type and were formed as a result of the Indian Plate colliding with the Eurasian
plate. Similarly the Appalachian Mountains of North America and the Alps of
Europe were formed by such processes.
Continental Rifting occurs where continental crust is undergoing extensional
deformation. This results in thinning of the lithosphere and upwelling of the
asthenosphere which results in uplift. The brittle lithosphere responds by
producing normal faults where blocks of continental lithosphere are uplifted to
form grabens or half grabens. The uplifted blocks are referred to a fault-block
The Basin and Range province in the western United states formed in this
manner, including the Sierra Nevada on its western edge and the Grand Tetons in
Cratons and Orogens
The continents can be divided into two kinds of structural units
Cratons form the cores of the continents. These are portions of continental crust
that have attained isostatic and tectonic stability and have cooled substantially
since their formation. They were formed and were deformed more than a billion
years ago and are the oldest parts of the continents. The represent the deep roots of
former mountains and consist of metamorphic and plutonic igneous rocks, all
showing extensive evidence of deformation.
Orogens are broad elongated belts of deformed rocks that are draped around the
cratons. They appear to be the eroded roots of former mountain belts that formed
by continent - continent collisions. Only the youngest of these orogens still form
mountain ranges (see figure 13.10) in your text).
The observation that the orogens are generally younger towards the outside of any
continent suggests that the continents were built by collisions of plates that added younger
material to the outside edges of the continents, and is further evidence that plate tectonics
has operated for at least the last 2 billion years.
Map Projections
The ways in which we visualize the world are varied- we have pictures, maps, globes,
satellite imagery, hand drawn creations and more. What kinds of things can we learn from
the way we see the world around us? For centuries mankind has been making maps of the
world around them, from their immediate area to the greater world as they understood it at
the time. These maps depict everything from hunting grounds to religious beliefs and
speculations of the broader, unexplored world around them. Maps have been made of the
local waterways, trade routes, and the stars to help navigators on land and sea make their
way to different locations. How we visualize the world not only has practical implications,
but can also help shape our perspectives of the Earth we live in.
There are many kinds of maps made from a variety of materials and on a variety of topics.
Clay tablets, papyrus, and bricks made way for modern maps portrayed on globes and on
paper; more recent technological advances allow for satellite imagery and computerized
models of the Earth. Certain map projections, or ways of displaying the Earth in the most
accurate ways by scale, are more well-known and used than other kinds. Three of these
common types of map projections are cylindrical, conic, and azimuthal.
Cylindrical Map Projections
Cylindrical map projections are one way of portraying the Earth. This kind of map
projection has straight coordinate lines with horizontal parallels crossing meridians at right
angles. All meridians are equally spaced and the scale is consistent along each parallel.
Cylindrical map projections are rectangles, but are called cylindrical because they can be
rolled up and their edges mapped in a tube, or cylinder. The only factor that distinguishes
different cylindrical map projections from one another is the scale used when spacing the
parallel lines on the map.
The downsides of cylindrical map projections are that they are severely distorted at the
poles. While the areas near the Equator are the most likely to be accurate compared to the
actual Earth, the parallels and meridians being straight lines don’t allow for the curvature
of the Earth to be taken into consideration. Cylindrical map projections are great for
comparing latitudes to each other and are useful for teaching and visualizing the world as
a whole, but really aren’t the most accurate way of visualizing how the world really looks
in its entirety.
Types of cylindrical map projections you may know include the popular Mercator
projection, Cassini, Gauss-Kruger, Miller, Behrmann, Hobo-Dyer, and Gall-Peters.
Conic Map Projections
Secondly, conic map projections include the equidistant conic projection, the Lambert
conformal conic, and Albers conic. These maps are defined by the cone constant, which
dictates the angular distance between meridians. These meridians are equidistant and
straight lines which converge in locations along the projection regardless of if there’s a
pole or not. Like the cylindrical projection, conic map projections have parallels that cross
the meridians at right angles with a constant measure of distortion throughout. Conic map
projections are designed to be able to be wrapped around a cone on top of a sphere
(globe), but aren’t supposed to be geometrically accurate.
Conic map projections are best suited for use as regional or hemispheric maps, but rarely
for a complete world map. The distortion in a conic map makes it inappropriate for use as
a visual of the entire Earth but does make it great for use visualizing temperate regions,
weather maps, climate projections, and more.
Azimuthal Map Projection
The azimuthal map projection is angular- given three points on a map (A, B, and C) the
azimuth from Point B to Point C dictates the angle someone would have to look or travel
in order to get to A. These angular relationships are more commonly known as great circle
arcs or geodesic arcs. The main features of azimuthal map projections are straight
meridian lines, radiating out from a central point, parallels that are circular around the
central point, and equidistant parallel spacing. Light paths in three different categories
(orthographic, stereographic, and gnomonic) can also be used. Azimuthal maps are
beneficial for finding direction from any point on the Earth using the central point as a
Map projection types all have their pros and cons, but they are incredibly versatile. Even
though it is nearly impossible to create an entirely accurate map projection there are uses
for even the most imperfect depictions of the Earth. Map projections are created for
certain purposes and should be used for those purposes. In the end each and every map
projection has a place, and there is no limit to the amount of projections that can be
Converting a sphere to a flat surface results in distortion. This is the most profound single
fact about map
projections—they distort the
world—a fact that you will
investigate in more detail in
Module 4, Understanding
and Controlling Distortion.
Imagine a map projection as
an attempt to reconstruct your
face in two dimensions.
Some maps will get the
shapes of all your features
just right, but not the
sizes—your forehead and
chin, for instance, may come
out huge. Other maps will get
the sizes right, but the shapes
will be stretched—maybe
your full, round mouth will
appear wide, thin, and rather
Some maps preserve distances. Measurements from the tip of your nose to your chin, ears,
and eyes will be right, even though the size and shape of your features is wrong. Other
maps preserve direction. Your features may look weird, and they may be scrunched up or
set too far apart, but their relative positions will be correct.
Finally, some maps are compromises—they get nothing exactly right but nothing too far
wrong. In particular, compromise projections try to balance shape and area distortion.
So the four spatial properties subject to distortion in a projection are:
If a map preserves shape, then feature outlines (like country boundaries) look the same on
the map as they do on the earth. A map that preserves shape is conformal. Even on a
conformal map, shapes are a bit distorted for very large areas, like continents.
A conformal map distorts area—most features are depicted too large or too small. The
amount of distortion, however, is regular along some lines in the map. For example, it may
be constant along any given parallel. This would mean that features lying on the 20th
parallel are equally distorted, features on the 40th parallel are equally distorted (but
differently from those on the 20th parallel), and so on.
If a map preserves area, then the size of a feature on a map is the same relative to its size
on the earth. For example, on an equal-area world map, Norway takes up the same
percentage of map space that actual Norway takes up on the earth.
To look at it another way, a coin moved to different spots on the map represents the same
amount of actual ground no matter where you put it.
In an equal-area map, the shapes of most features are distorted. No map can preserve both
shape and area for the whole world, although some come close over sizeable regions.
If a line from a to b on a map is the same distance (accounting for scale) that it is on the
earth, then the map line has true scale. No map has true scale everywhere, but most maps
have at least one or two lines of true scale.
An equidistant map is one that preserves true scale for all straight lines passing through a
single, specified point. For example, in an equidistant map centered on Redlands,
California, a linear measurement from Redlands to any other point on the map would be
Direction, or azimuth, is measured in degrees of angle from north. On the earth, this means
that the direction from a to b is the angle between the meridian on which a lies and the
great circle arc connecting a to b.
The azimuth of a to b is 22 degrees.
If the azimuth value from a to b is the same on a map as on the earth, then the map
preserves direction from a to b. An azimuthal projection is one that preserves direction for
all straight lines passing through a single, specified point. No map has true direction
A few projections with different properties. The Lambert Conformal Conic preserves
shape. The Mollweide preserves area. (Compare the relative sizes of Greenland and South
America in one and then the other.) The Orthographic projection preserves direction. The
Azimuthal Equidistant preserves both distance and direction. The Winkel Tripel is a
compromise projection.
More about scale
Scale is the relationship between distance on a map or globe and distance on the earth.
Suppose you have a globe that is 40 million times smaller than the earth. Its scale is
1:40,000,000. Any line you measure on this globe—no matter how long or in which
direction—will be one forty-millionth as long as the corresponding line on the earth. In
other words, the scale is true everywhere. This is because the globe and the earth have the
same shape (disregarding the complication of sphere versus spheroid).
Now suppose you have a flat map that is 40 million times smaller than the earth. (See the
problem coming? Instead of comparing a big orange to a little orange, we're comparing a
big orange to a little wafer.) This map also has a scale of 1:40,000,000, but because the
map and the earth are differently shaped, this scale cannot be true for every line on the
The stated scale of a map is true for certain lines only. Which lines these are depends on
the projection and even on particular settings within a projection. We'll come back to this
subject in Module 4, Understanding and Controlling Distortion.
Not all of the earth's curves can be represented as straight lines at the same fixed scale.
Some lines must be shortened (and others lengthened).
Expressing map scale
There are three common ways to express map scale:
Linear scales
Linear scales are lines or bars drawn on a map with real-world distances marked on them.
To determine the real-world size of a map feature, you measure it on the map with a ruler
or a piece of string. Then you compare the feature's length on the string to the scale bar.
A typical scale bar.
Verbal scales
Verbal scales are statements of equivalent distances. For example, if a 4.8 kilometer road
is drawn as a 20 centimeter line on a map, a verbal scale would be “20cm = 4.8km.” You
could also formulate the scale (reducing both sides by 20) as “1cm = .24km.”
Representative fractions
Representative fractions express scale as a fraction or ratio of map distance to ground
distance. For example, a scale of 1:24,000 (also written 1/24,000) means that one unit on
the map is equal to 24,000 of the same units on the earth. Since the scale is a ratio, it
doesn't matter what the units are. You can interpret it as 1 meter = 24,000 meters, 1 mile =
24,000 miles, or 1 hand width = 24,000 hand widths (as long as it's the same hand).
Small scale and large scale maps
It's easy to mix these terms up. Here's one way to keep them straight: on a large-scale map,
the earth is large (so not very much of it fits on the map). On a small-scale map, the earth
is small (so all or most of it fits on the map). A map of your town, or your property, is
going to be a large-scale map. A continental or world map is a small-scale map.
Another way to think of the difference in terms of representative fractions. The larger the
fraction, the larger the map's scale. For example, 1/10,000 is a larger fraction than
1/1,000,000. So a 1:10,000 map is larger scale than a 1:1,000,000 map.
Measuring distortion using Tissot's Indicatrix
In the nineteenth century, Nicolas Auguste Tissot developed a method to analyze map
projection distortion.
An infinitely small circle on the earth's surface will be projected as an infinitely small
ellipse on any given map projection. The resulting ellipse of distortion, or indicatrix,
shows the amount and type of distortion at the location of the ellipse.
For example, if an indicatrix is elongated from north to south, shape is correspondingly
distorted at that location on the map. The same goes for east–west stretching or oblique
stretching. On a conformal map, the indicatrices are all circles, but they vary in size. On an
equal area projection, the indicatrices have varying ellipticity, but the same area.
A cylindrical projection is any projection in which the meridians are mapped to parallel
spaced vertical lines and latitudes are mapped to horizontal lines. The projections stretch
from east to west according to their geometric constructions and are the same at any
chosen latitude. Cylindrical projections are distinguished from each other by the north to
south stretching denoted by φ. The north to south stretching equals east to west but grows
with latitude faster than east to west stretching in the case of central cylindrical projection.
Mercator projection is an example of cylindrical projection which became a standard map
projection because of its ability to represent lines of steady course. Mercator distorts the
size of geographical objects because its linear scale increases with the increase in latitude.
The distortion caused by the Mercator distorts the perception of the entire planet by
exaggerating the areas laying far from the equator.
Pseudocylindrical projections present the meridian as a straight line while other parallels
as sinusoidal curves which are longer than the central meridian. The scaling of the
pseudocylindrical projections are straight along the central meridian and also along the
parallels. On a pseudocylindrical map, points further from the equator have higher
latitudes than other points, preserving the north-south relationship. Pseudocylindrical
projections include sinusoidal with same horizontal and vertical scales. The Robinson
projection was created to show the globe as a flat image readily. The projection is neither
equal-area nor conformal because of the compromise to show the whole planet. The
meridians of the Robinson projection curves are gently stretching the poles into long lines.
Van der Grinten Projection
Van der Grinten is a compromised projection which is neither equal-area nor conformal. It
is an arbitrary scaled projection of the plane projecting the entire earth into a circle. Van
der Grinten projection preserves the image of Mercator projection and reduces its
distortion. However, the Polar Regions can still be distorted by the Van der Grinten
Conic Projection
Conic projections have meridians mapped to equally spaced parallels originating from the
top while the parallels are mapped to circular arcs which are centered at the top. Two
standard lines visualized as secant lines are picked in the process of making a conic
projection. When a single parallel line has used the distance along the parallels is
stretched. Examples of conic maps include equidistant, Albers, and Lambert conformal
Pseudoconic Projection
Pseudoconic Projections are projections with parallels which are circular arcs with
common central points. Unlike conic projections, the meridian is not constrained to be a
straight line. Examples of pseudoconic projections include "bonne", which is an
equal-area map projection. The maps are not constrained to rectangles or discs.
Pseudoconic projection is one of the oldest map types and although they were used by
Ptolemy, they are seldom seen today.
Scale Distortions On Map Projections
Map projections without distortions would represent the correct distance, direction,
shapes, and areas on a map. However, map projections have distortions which depend
largely on the size of the area being mapped. Scale distortions on maps are shown on the
map by an ellipse of distortion or using scale factor which is the ratio of the scale at a
given point to the true scale. Distortions on maps of countries or cities are not evident to
the eye and can only be identified when computing distances and areas.
Characteris cs of Map Projec ons
The general purpose of map projections and the basic problems encountered have been discussed often
and well in various books on cartography and map projections. (Robinson, Sale, Morrison, and
Muehrcke, 1984; Steers, 1970; and Greenhood, 1964, are among later editions of earlier standard
references.) Every map user and maker should have a basic understanding of projections, no matter
how much computers seem to have automated the operations. The concepts will be concisely described
here, although there are some interpretations and forrnulas that appear to be unique.
For almost 500 years, it has been conclusively established that the Earth is essentially a sphere,
although a number of intellectuals nearly 2,000 years earlier were convinced of this. Even to the
scholars who considered the Earth flat, the skies appeared hemispherical, however. It was established at
an early date that attempts to prepare a flat map of a surface curving in all directions leads to distortion
of one form or another.
A map projection is a systematic representation of all or part of the surface of a round body, especially
the Earth, on a plane. This usually includes lines delineating meridians and parallels, as required by
some definitions of a map projection, but it may not, depending on the purpose of the map. A projection
is required in any case. Since this cannot be done without distortion, the cartographer must choose the
characteristic which is to be shown accurately at the expense of others, or a compromise of several
characteristics. If the map covers a continent or the Earth, distortion will be visually apparent. If the
region is the sizc of a small town, distortion may be barely measurable using many projections, but it
can still be serious with other projections. There is literally an infinite number of map projections that
can be devised, and several hundred have been published, most of which are rarely used novelties. Most
projections may be infinitely varied by choosing different points on the Earth as the center or as a
starting point. It cannot be said that there is one "best" projection for mapping. It is even risky to claim
that one has found the "best" projection for a given application, unless the parameters chosen are
artificially constricting. A carefully constructed globe is not the best map for most applications because
its scale is by necessity too small. A globe is awkward to use in general, and a straightedge cannot be
satisfactorily used on one for measurement of distance.
The details of projections discussed in this book are based on perfect plotting onto completely stable
media. In practice, of course, this cannot be achieved. The cartographer may have made small errors,
especially in hand-drawn maps, hut a more serious problem results from the fact that maps are
commonly plotted and printed on paper, which is dimensionally unstable. Typical map paper can
expand over 1 percent with a 60 percent increase in atmospheric humidity, and the expansion
coefficicnt varies considcrably in different directions on the same sheet. This is much greater than the
variation between common projections on largescale quadrangles, for example. The use of stable plastic
bases for maps is recommended for precision work, but this is not always feasible, and source maps may
be available only on paper, frequently folded as well. On large-scale maps, such as topographic
quadrangles, measurement on paper maps is facilitated with rectangular grid overprints, which expand
with the paper. Grids are discussed later in this book.
The characteristics normally considered in choosing a map projection are as follows:
Area - Many map projections are designed to be equal-area, so that a coin of any size, for
example, on one part of the map covers exactly the same area of the actual Earth as the same
coin on any other part of the map. Shapes, angles, and scale must be distorted on most parts of
such a map, but there are usually some parts of an equal-area map which are designed to retain
these characteristics correctly, or very nearly so. Less common terms used for equal-area
projections are equivalent, homolographic, or homalographic (from the Greek homalos or
homos ("same") and graphos ("write")); authalic (from the Greek autos ("same") and ailos
("area")), and equiareal.
Shape - Many of the most common and most important projections are conformal or
orthomorphic (from the Greek orthos or "straight" and morphe or "shape"), in that normally
the relative local angles about every point on the map are shown correctly. (On a conformal
map of the entire Earth there are usually one or more "singular" points at which local angles are
still distorted.) Although a large area must still be shown distorted in shape, its small features
are shaped essentially correctly. Conformality applies on a point or infinitesimal basis, whereas
an equal-area map projection shows areas correctly on a finite, in fact mapwide basis. An
important result of conformality is that the local scale in every direction around any one point
is constant. Because local angles are correct, meridians intersect parallels at right (90�) angles
on a conformal projection, just as they do on the Earth. Areas are generally enlarged or reduced
throughout the map, but they are correct along certain lines, depending on the projection.
Nearly all large-scale maps of the Geological Survey and other mapping agencies throughout
the world are now prepared on a conformal projection. No map can be both equal-area and
While some have used the term aphylactic for all projections which are neither equal-area nor
conformal (Lee, 1944), other terms have commonly been used to describe special
Scale - No map projection shows scale correctly throughout the map, but there are usually one
or more lines on the map along which the scale remains true. By choosing the locations of these
lines properly, the scale errors elsewhere may be minimized, although some errors may still be
large, depending on the size of the area being mapped and the projection. Some projections
show true scale between one or two points and every other point on the map, or along every
meridian. They are called equidistant projections.
Direction - While conformal maps give the relative local directions correctly at any given
point, there is one frequently used group of map projections, called azimuthal (or zenithal), on
which the directions or azimuths of all points on the map are shown correctly with respect to
the center. One of these projections is also equal-area, another is conformal, and another is
equidistant. There are also projections on which directions from two points are correct, or on
which directions from all points to one or two selected points are correct, but these are rarely
Special characteristics - Several map projections provide special characteristics that no
other projection provides. On the Mercator projection, all rhumb lines, or lines of constant
direction, are shown as straight lines. On the Gnomonic projection, all great circle paths-the
shortest routes between points on a sphere shown as straight lines. On the Stereographic, all
small circles, as well as great circles, are shown as circles on the map. Some newer projections
are specially designed for satellite mapping. Less useful but mathematically intriguing
projections have been designed to fit the sphere conformally into a square, an ellipse, a triangle,
or some other geometric figure.
Method of construction - In the days before ready access to computers and plotters, ease of
construction was of greater importance. With the advent of computers and even pocket
calculators, very complicated formulas can be handled almost as routinely as simple projections
in the past.
While the above six characteristics should ordinarily be considered in choosing a map projection, they
are not so obvious in recognizing a projection. In fact, if the region shown on a map is not much larger
than the United States, for example, even a trained eye cannot often distinguish whether the map is
equal-area or conformal. It is necessary to make measurements to detect small differences in spacing or
location of meridians and parallels, or to make other tests. The type of construction of the map
projection is more easily recognized with experience, if the projection falls into one of the common
There are three types of developable1 surfaces onto which most of the map projections used by the
USGS are at least partially geometrically projected. They are the cylinder, the cone, and the plane.
Actually all three are variations of the cone. A cylinder is a limiting form of a cone with an increasingly
sharp point or apex. As the cone becomes flatter, its limit is a plane. If a cylinder is wrapped around the
globe representing the Earth (see fig. l), so that its surface touches the Equator throughout its
circumference, the meridians of longitude may be projected onto the cylinder as equidistant straight
lines perpendicular to the Equator, and the parallels of latitude marked as lines parallel to the Equator,
around the circumference of the cylinder and mathematically spaced for certain characteristics. For
some cases, the parallels may also be projected geometrically from a common point onto the cylinder,
but in the most common cases they are not perspective. When the cylinder is cut along some meridian
and unrolled, a cylindrical projection with straight meridians and straight parallels results. The
Mercator projection is the best-known example, and its parallels must be mathematically spaced.
If a cone is placed over the globe, with its peak or apex along the polar axis of the Earth and with the
surface of the cone touching the globe along some particular parallel of latitude, a conic (or conical)
projection can be produced. This time the meridians are projected onto the cone as equidistant straight
lines radiating from the apex, and the parallels are marked as lines around the circumference of the
cone in planes perpendicular to the Earth's axis, spaced for the desired characteristics. The parallels
may not be projected geometrically for any useful conic projections. When the cone is cut along a
meridian, unrolled, and laid flat, the meridians remain straight radiating lines, but the parallels are now
circular arcs centered on the apex. The angles between meridians are shown smaller than the true
A plane tangent to one of the Earth's poles is the basis for polar azimuthal projections. In this case, the
group of projections is named for the function, not the plane, since all common tangent-plane
projections of the sphere are azimuthal. The meridians are projected as straight lines radiating from a
point, but they are spaced at their true angles instead of the smaller angles of the conic projections. The
parallels of latitude are complete circles, centered on the pole. On some important azimuthal
projections, such as the Stereographic (for the sphere), the parallels are geometrically projected from a
common point of perspective; on others, such as the Azimuthal Equidistant, they are nonperspective.
FIGURE 1.-Projection of the Earth onto the three major surfaces. In a few cases, projection is geometric,
but in most cases the projection is mathematical to achieve certain features.
The concepts outlined above may be modified in two ways, which still provide cylindrical, conic, or
azimuthal projections (although the azimuthals retain this property precisely only for the sphere).
The cylinder or cone may be secant to or cut the globe at two parallels instead of being tangent
to just one. This conceptually provides two standard parallels; but for most conic projections
this construction is not geometrically correct. The plane may likewise cut through the globe at
any parallel instead of touching a pole, but this is only useful for the Stereographic and some
other perspective projections.
The axis of the cylinder or cone can have a direction different from that of the Earth's axis,
while the plane may be tangent to a point other than a pole (fig. 1). This type of modification
leads to important oblique, transverse, and equatorial projections, in which most meridians
and parallels are no longer straight lines or arcs of circles. What were standard parallels in the
normal orientation now become standard lines not following parallels of latitude.
Other projections resemble one or another of these categories only in some respects. There are
numerous interesting pseudocylindrical (or "false cylindrical") projections. They are so called because
latitude lines are straight and parallel, and meridians are equally spaced, as on cylindrical projections,
but all meridians except the central meridian are curved instead of straight. The Sinusoidal is a
frequently used example. Pseudoconic projections have concentric circular arcs for parallels, like conics,
but meridians are curved; the Bonne is the only common example. Pseudoazimuthal projections are
very rare; the polar aspect has concentric circular arcs for parallels, and curved meridians. The
Polyconic projection is projected onto cones tangent to each parallel of latitude, so the meridians are
curved, not straight. Still others are more remotely related to cylindrical, conic, or azimuthal
projections, if at all.
Stereonet Projections
The stereographic projection is a methodology used in structural geology and engineering
to analyze orientation of lines and planes with respect to each other. The stereonets is a
type of standardized mapping system that allows us to represent various angles in 3D
space on a 1D paper. They are used for analysis of various field data such as bedding
attitudes, planes, hinge lines and numerous other structures. This is a very useful tool
because it can reduce the workload by avoiding lengthy calculations.
In structural geology, we use the bottom half or hemisphere of the spherical projection. If
you are a mineralogist, you will use the top half of the spherical projection for
crystallographic analysis. The reasoning behind which hemisphere we used is more
conceptual than anything. This will be explained in depth in a different article. What is
important to someone who just started using steronets is to recognize that steronets
represents half a sphere where the cross section has 360-degrees. The pole to the plane
(“dip pole”) is at 90-degrees to the plane. Planes are lines are drawn on steronets as they
intersect at the bottom of the sphere (Figure 1).
Figure 1: Steronet with a plane and its pole. Note the bottom half of a sphere is used.
Click to enlarge.
Types of Steronets
There are two widely used types (and may be more) of stereonets by structural geologists.
They are equal area stereonet and equal angle stereonets. The choice either should not
affect the data analysis. The analysis and interpretation of data achieved through the use of
either equal area of equal angle steronets should result in same conclusions.
However, the equal area steronets will reduce the area distortion. In other words, it is often
used to analyze accuracy of data from several different regions of the same area. It is also
useful in structure stereonet contouring. Hence, most educational institutions prefer equal
area steronets for their students over the equal angle stereonets.
The equal angle stereonets are suitable for kinematic analysis. In other words, they
provide the best projection for analyzing the direction and the vectors of structural forces.
This is because the equal angle stereonets preserves the true relationships between
stratigraphic and structural features.
For our discussion here, I will be using Wulff-Lambert type, which will preserve the
North:It is the true North which is denoted by the azimuthal angle of 000-degrees on the
primitive. All strike angles are measured with respect to the true North.
Primitive: It is the outer most circle is the primitive. It is at 90-degrees from the center of
the stereonet. Primitive circle is also a great circle but, it contains N, E, S and W directions
at 000, 090, 180 and 270 degrees intervals.
Great circle: A circle on the surface of the sphere made by the intersection with the
spehere of a plane that passes through the center of the sphere. The great circles run
North-South (longitudinal) or up-down and bisect the sphere precisely. The great circle is
divided in to 360 degrees (like 360 degree protractor) because maps are designed based on
same azimuthal (bearing) directional vectors. If you have understand how 3D vectors
work, this should be a no-brainer.
Small circle: A circle on the surface of a sphere made by the intersections of a plane that
does not pass through the center of the sphere. Small circles run left-right (latitudinal) on
the stereonets and are perpendicular to the great circles.
Figure 2: Highlighted in green is the primitive. The blue arrows are pointing to two great
circle. The orange arrows are pointing to two small circles. Click to enlarge.
Types of Geological Plotting and Data Usage
– Bedding surface, fault or a structure (“features”): strike of any of these will plot as a line
on the stereonet.
● strike measured using the great circle
● always rotate the tick mark for the strike to the North, then counts from East
(right hand side) towards the center for dip (or use the right hand rule)
● use same principles for overturned beds (you cannot differentiate an overturned
bed from a regular bed on a steronet)
– Pole to a feature: is always exactly 90-degrees opposite to the dip direction of the
● should me marked as a dot with a circle around it
– Trend: tread is always taken on line to a plane (NOT on the line)
● usually deal with folds axis or a fault (ONLY on intersection of the line)
● move the intersection to the North (or to a straight line) and make a tick/line on
the great circle where the intersection meets
● move back the trace paper to align with North; the point at which the tick
intersect the great circle is the trend
– Plunge: it provides the angular information on how deep a structure is dipping in to a
surface; very common feature in coal beds and folds
● plunge is the distance between the great circle and the intersection
– Rake (pitch): distance between the intersection and the great circle along a plane/line
(must be less than 90) rake is the angle between strike and trend
Normal fault with the Net-slip displacement and horizontal vector.
On the animation above, I drew two vectors out of several which can be used to interpret a
normal fault. The red arrow is the displacement vector which is obtained by the horizontal
and vertical displacement. The horizontal displacement is indicated with the brown arrow
(vertical displacement is NOT shown). The rake of the fault is between the left most edge
of the footwall and the displacement vector(red). The green arrow represents the rate of
drop with respect to the original block. We use slickensides to interpret the sense of
motion in the field.
Example with Data
Plane A: 040/36, Plane B: 320/40, Pi-girdle to Plane A and B (line of best fit crossing
points 1 and 2), the hinge at point 3
Based on the above diagram…
-There are two planes; A and B
-There are two rake angles measured on the planes itself; measured between the
intersection point 3 and the great circle. For example, from intersection point 3 upwards
towards NW direction of the great circle intersection of plane A. Plane B rake is
downwards towards SE direction.
-The pi-gridle is determined by plotting poles to the Plain A and B at 90 degrees, then
alining them on one of the lines on the stereonet. The point 1 and 2 are best fit line points
for the poles that lies about the center of the diagram.
-Trend is the along the point 1 and 3, directly outwards on to the great circle (it is NOT
marked in this diagram). It is measured on the great circle itself.
-Plunge is the distance between the great circle-trend intersection and the point 3.
A detailed diagram…
Hand written sample
Manual versus Software
There are absolutely no
differences between the
interpretations made using
manual drawing and
software-based drawing of
datasets. In work environment,
we usually use software to
generate stereonets. The
software often eliminates many
user errors, produce much better
quality steronets extremely
detailed analysis of datasets and
make it easier to share with
other over electronic devices.
For someone who is starting in
geology or structural geology, it is highly recommended to use paper and pencil over
software. This will help you learn the fundamentals of stereographic projection. Typically
university geology and engineering students are expected create stereonets by hand.
Geohazard Risks:
Geology (from the Greek γῆ, gê, "earth" and λόγος, logos, "study") is the science
comprising the study of solid Earth, the rocks of which it is composed, and the processes
by which it evolves.
Hazard: In the context of PanGeo, is any geologically-related phenomena that has
potential to cause harm.
In PanGeo only geohazards related to ground instability are included, thereby ignoring
geohazards relating to radon gas or flooding.
Geohazards fall into various categories as follows (described in more detail below):
Earthquake (seismic) hazard
Tectonic movements
Salt Tectonics
Volcanic Inflation/deflation
Soil Creep
Ground Dissolution
Collapsible Ground
Running Sand/ Liquefaction
● Shrink-swell clays
● Compressible Ground
Ground water management - Shallow compaction
Ground water management - Peat oxidation
Groundwater abstraction
● Underground construction
● Made ground
● Oil and Gas Production
● Earthquake (seismic) hazard: Earthquakes are the observable effects of vibrations
(known as seismic waves) within the Earth’s crust arising from relatively rapid
stress release, typically along a fault zone. Damage to buildings and other
infrastructure can be caused as the ground shakes during the passage of seismic
waves. Other effects include liquefaction of water-saturated soft ground,
potentially leading to a loss in ground strength and the extrusion of water-saturated
sediments as ‘mud volcanoes’ and the like. Ground shaking can also trigger
secondary events such as landslides and tsunami. Some earthquakes are associated
with significant permanent vertical or lateral ground movement. Changes to
drainage systems can cause flooding. There is potential for injury and loss of life
during earthquakes. Seismic hazard can be assessed by reference to the size and
frequency of recorded earthquakes, although individual earthquakes are essentially
unpredictable. Individual events occur on time-scales of seconds or minutes.
Modern infrastructure should be designed to withstand probable local seismic
● Tectonic movements: Tectonic movements are large scale processes that affect the
Earth’s crust. These processes can lead to areas of the crust rising or falling.
Neotectonic movements are typically due to the stresses introduced through
movements of the Earth’s plates. These types of motion are likely to be on a broad
scale and so it may not be possible to measure them using satellite radar.
● Salt Tectonics: Localised motions can be associated with the movement of
evaporate deposits, these are termed salt tectonics and can produce both uplift and
subsidence depending on the exact mechanisms at play.
● Volcanic Inflation/deflation: Volcanic activity can lead to the creation of lava
flows, ash flows, debris and ash falls, and debris flows of various kinds. It might
be accompanied by release of poisonous or suffocating gases, in some instances
with explosive violence, or by significant seismic activity or ground movement.
Secondary effects can include landslide and flooding. For PanGeo we are
interested in hazards associated with ground instability. Ground instability
associated with volcanoes tends to relate to inflation and deflation of the ground
surface as magma volumes change. Secondary effects such as landslides should be
mapped into the other relevant PanGeo geohazard classes.
The propensity for upward, lateral or downward movement of the ground can be caused
by a number of natural geological processes. Some movements associated with particular
hazards may be gradual or occur suddenly and also may vary from millimetre to metre or
tens of metres scale. Note that anthropogenic deposits can be affected by natural ground
instability. Significant natural ground instability has the potential to cause damage to
buildings and structures, and weaker structures are most likely to be affected. It should be
noted, however, that many buildings, particularly more modern ones, are built to such a
standard that they can remain unaffected in areas of even significant ground movement.
The susceptibility of built structures to damage from geohazards might also depend on
local factors such as the type of nearby vegetation, or the nature of the landforms in the
area. The effects of natural ground instability often occur over a local area as opposed to
the effects of natural ground movements which occur over larger areas.
● Landslide: A landslide is a relatively rapid outward and downward movement of
a mass of rock or soil on a slope, due to the force of gravity. The stability of a
slope can be reduced by removing ground at the base of the slope, increasing the
water content of the materials forming the slope or by placing material on the
slope, especially at the top. Property damage by landslide can occur through the
removal of supporting ground from under the property or by the movement of
material onto the property. Large landslides in coastal areas can cause tsunami.
The assessment of landslide hazard refers to the stability of the present land
surface, including existing anthropogenically-modified slopes as expressed in local
topographic maps or digital terrain models. It does not encompass a consideration
of the stability of new excavations. Land prone to landslide will normally remain
stable unless the topography is altered by erosion or excavation, or the land is
loaded, or pore water pressure increases. Landslide might also be initiated by
seismic shock, frost action, or change in atmospheric pressure. This hazard is
significant in surface deposits but may extend to more than 10m depth. The
common consequences are damage to properties, including transportation routes
and other kinds of infrastructure, and underground services. Some landslides can
be stabilised by engineering.
● Soil Creep: Soil creep is a slow movement of soil and rock particles down-slope
and is a result of expansion and contraction of the soil through cycles of freezing
and thawing or wetting and drying.
● Ground Dissolution: Some rocks and minerals are soluble in water and can be
progressively removed by the flow of water through the ground. This process
tends to create cavities, potentially leading to the collapse of overlying materials
and possibly subsidence at the surface. The common types of soluble rocks and
minerals are limestones, gypsum and halite. Cavities can become unstable
following flooding, including flooding caused by broken service pipes. Changes in
the nature of surface runoff, excavating or loading the ground, groundwater
abstraction, and inappropriate installation of soakaways can also trigger subsidence
in otherwise stable areas.
● Collapsible Ground: Collapsible ground comprises materials with large spaces
between solid particles. They can collapse when they become saturated by water
and a building (or other structure) places too great a load on it. If the material
below a building collapses it may cause the building to sink. If the collapsible
ground is variable in thickness or distribution, different parts of the building may
sink by different amounts, possibly causing tilting, cracking or distortion. Collapse
will occur only following saturation by water and/or loading beyond criticality.
This hazard can be significant in surface deposits and possibly also in buried
superficial deposits.
● Running Sand/ Liquefaction: Running sand occurs when loosely-packed sand,
saturated with water, flows into an excavation, borehole or other type of void. The
pressure of the water filling the spaces between the sand grains reduces the contact
between the grains and they are carried along by the flow. This can lead to
subsidence of the surrounding ground. If sand below a building runs it may
remove support and the building may sink. Different parts of the building may
sink by different amounts, possibly causing tilting, cracking or distortion. The
common consequences are damage to properties or underground services. This
hazard tends to be self-limited by decrease in head of water.
Liquefaction of water-saturated soft ground often results as an effect of earthquake
activity but can also be triggered by manmade vibrations due to construction works. It
can potentially lead to a loss in ground strength and the extrusion of water-saturated
sediments as ‘mud volcanoes’ and the like. Soils vulnerable to liquefaction represent
areas of potential ground instability.
● Shrink-swell clays: A shrinking and swelling clay changes volume significantly
according to how much water it contains. All clay deposits change volume as their
water content varies, typically swelling in winter and shrinking in summer, but
some do so to a greater extent than others. Most foundations are designed and built
to withstand seasonal changes. However, in some circumstances, buildings
constructed on clay that is particularly prone to swelling and shrinking behaviour
may experience problems. Contributory circumstances could include drought,
leaking service pipes, tree roots drying-out of the ground, or changes to local
drainage such as the creation of soakaways. Shrinkage may remove support from
the foundations of a building, whereas clay expansion may lead to uplift (heave) or
lateral stress on part or all of a structure; any such movements may cause cracking
and distortion. The existence of this hazard depends on a change in soil moisture
and on differential ground movement. Uniform ground movement may not of
itself present a hazard. This hazard is generally significant only in the top five
metres of ground.
● Compressible Ground: Many ground materials contain water-filled pores (the
spaces between solid particles). Ground is compressible if a building (or other
load) can cause the water in the pore space to be squeezed out, causing the ground
to decrease in thickness. If ground is extremely compressible the building may
sink. If the ground is not uniformly compressible, different parts of the building
may sink by different amounts, possibly causing tilting, cracking or distortion.
This hazard commonly depends on differential compaction, as uniform compaction
may not of itself present a hazard. Differential compaction requires that some
structure that might be susceptible to subsidence damage has been built on
non-uniform ground. The common consequences are damage to existing
properties that were not built to a sufficient standard, and possible damage to
underground services.
These are ground motions covering a local area which have been brought about by the
activity of man. Subsidence (downward movement) of the ground can result from a
number of different types of anthropogenic activity, namely mining (for a variety of
commodities), or tunnelling (for transport, underground service conduits, or for
underground living or storage space). Subsidence over a regional area can result from
fluid abstraction (for water, brine, or hydrocarbons). Uplift or heave of the ground can
occur when fluid is allowed to move back into an area from where it was previously
extracted and groundwater recharge occurs. This fluid recovery may include injection of
water or gas.
● Ground water management - Shallow compaction: Ground water management
may be applied for example to ensure the exploitability of existing agricultural
land in lowland coastal areas. Groundwater management can lead to higher or
lower water levels of phreatic groundwater and of deeper aquifers in the shallow
subsurface. Groundwater occupies pore and interstitial spaces and fractures within
sediments and rocks and therefore exerts a pressure. When the water is drained the
pore pressure or effective stress is reduced. This leads to consolidation of
especially soft sediments, such as clay and peat. This change in the sediment
volume leads to subsidence. Similarly when groundwater levels are allowed to
recover, uplift may be a result of increasing pore pressure.
● Ground water management - Peat oxidation: Peat oxidation is the chemical
reaction where peat starts decomposing and will waste away with time. This loss
of soil volume leads to subsidence. It occurs when layers of peat in the subsurface
are exposed to oxygen. As long as peat is located in saturated ground layers this
process does not take place. However peat oxidation does occur in unsaturated
soils, for instance in areas where ground water management lowers ground water
Groundwater abstraction: Groundwater occupies pore and interstitial spaces and
fractures within sediments and rocks in the deeper subsurface. When this water is
removed, for instance through pumping for drinking water or lowering of water
levels in mines, the pore pressure or effective stress is reduced and consolidation of
the sediments causes a change in the sediment volume. This leads to subsidence.
Similarly when aquifer levels are allowed to recover, uplift may be a result of
increasing pore pressure. Deep geothermal energy systems should not lead to
ground movement. They involve closed systems where water, which was extracted
from a deep aquifer, will be pumped back into that same aquifer. However,
geothermal heat pumps are used at shallower depths. Although these are also
closed systems, ground movement might occur temporarily (e.g. seasonally) or
even permanently.
Mining: Mining is the removal of material from the ground, in the context of
PanGeo we consider mining to relate to the removal of solid minerals. The ground
surface may experience motion due to readjustments in the overburden if
underground mine workings fail.
Underground construction: In PanGeo we are interested in underground
construction that might bring about ground instability. An example of this would
be underground tunnelling; the removal of subsurface material can alter the
support for the overlying material therefore leading to ground motions.
Made ground: Made ground comprises anthropogenic deposits of all kinds such as
land reclamation, site and pad preparation by sand infill, road and rail
embankments, levees and landfills for waste disposal. Examples of land
reclamation are artificial islands, beach restoration and artificial harbours.
Reclaimed land as well as embankments and levees are generally made up of sand,
which is not prone to compaction as are clay and peat. However, two ground
instability processes will occur: consolidation of this artificial ground and
compaction of the ground below due to the load of the artificial ground and the
structure it supports, e.g. a building. Depending on its composition and mode of
deposition, landfill can also be a compressible deposit.
Oil and Gas Production: Similar to abstraction of groundwater the production of
oil and gas decreases the pore pressure of the reservoir rocks and therefore can
cause consolidation and subsidence of the surface. Storage of material in the
depleted reservoir (such as natural gas or CO2) can lead to surface uplift.
What is a geohazard?
Geohazards can be defined as events related to the geological state and processes
that may cause loss of lives as well as material and environmental damages. These
geohazards all arise from global geological processes inside the Earth, driving
deformation and displacement of its crust. Underneath the thin crust the Earth
consists of a sticky fluid of melted rock we call the mantle that turns and twists like
boiling water in slow motion, causing the crust to move ever so slowly.
Conceptual drawing of assumed convection cells in the mantle. Compared with boiling
of water.Credit: USGS
The crust is divided in different plates (tectonic plates) and when these plates interact
the resulting crustal movement can cause earthquakes, allow volcanoes to erupt and
set off landslides. All of these three; earthquakes, volcanoes and landslides can
trigger tsunamis if they happen in or close to the ocean. These four geohazards are
what I call the Fantastic Four in Planet Earth: Extreme Beauty – Extreme Danger.
Earthquakes: Fractures in Earth's crust, or lithosphere (its crust and upper mantle),
where sections of rock have slipped past each other are called faults. Earthquakes are
caused by the sudden release of accumulated strain along these faults, releasing
energy in the form of low-frequency sound waves called seismic waves. A major
earthquake are usually followed by aftershocks. The epicenters of large earthquakes
are normally located along known seismically active zones, although the disruptive
effects of an earthquake may extend over areas 100s of kilometers away.
Earthquakes may cause liquefaction, landslides, marine landslides and tsunamis.
Volcanoes: A volcano is defined by an opening in the Earth's crust from which lava,
ash, and hot gases flow or are ejected during an eruption. Volcanic hazards vary from
one volcano to another and from one eruption to the next. The big killers are
pyroclastic flows, lahars, and tsunamis triggered by volcanic eruptions. The most
frequent lethal events are so-called tephra explosions – very rapid jets of lava . The
longest-lasting damage is usually inflicted by thick lava flows or major collapses of
volcanic edifices, as at Mt. St. Helens in 1980.
Volcanoes are spectacular
sights. This is from Hawaii.
Landslides: A landslide is a
geohazard that involves the
breakup and downhill flow of
rock, mud, water and anything
caught in the path. Landslides
are one of the main processes
by which landscapes evolve and
the related hazards result in a
complex, changing landscape.
Landslides both vary enormously in their distribution in space and time, the amounts
of energy produced during the activity and especially in size. This means that the
resulting surface deformation or displacement varies considerably from one type of
instability (that trigger the breakup) to another. Individual ground instabilities may
have a common trigger, such as an extreme rainfall event or an earthquake, and
therefore occur alongside many equivalent occurrences over a large area. This means
that they can have a significant regional impact.
Landslide in Tibet. Clearing of roads is done with
manual labor. Photo: Bente Lilja Bye
Tsunamis: Tsunamis are gravity waves (different
physical features than wind induced surface waves)
created by a rapid displacement of a water column.
The displacement can be the result of earthquakes,
volcanic eruptions or landslides. These energetic
waves travel fast with long wavelengths and relatively
small amplitudes in open ocean. When hitting shallow
water they build up an amplitude and can become
tens, and on very rare occasions, even hundreds of
meters high. The coastal inundation can be devastating and catastrophic.
Extreme geohazards
The Sumatran earthquake/Indian ocean tsunami was
one of the most extreme geohazard in modern history
(see list of earthquakes). This extreme reached us all,
beyond the mere geophysical waves. As tourists come
from all over the world to visit the beautiful shores of
the Indian ocean, the 2004 tsunami affected people
from around the globe. The extent of it's destruction
and the dimension of the disaster are parts of the
definition of extreme. That, combined with their
physical features that normally are several orders
higher or more powerful than the average geohazard.
So, when we talk about extreme geohazards we not
only refer to the physical characteristics of the
geohazard but also the risk they represent in terms of consequences of this hazard.
A geohazard is a geological state that may lead to widespread damage or risk. Geohazards
are geological and environmental conditions and involve long-term or short-term
geological processes. Geohazards include earthquakes, volcanic activity, landslides,
tsunamis, etc. and can range from local events such as a rock slide or coastal erosion to
events that threaten humankind such as a supervolcano or meteorite impact. Earth
scientists undertake research to better understand these hazards and contribute to risk
management policies related to social and technical issues associated with geohazards as
well as disaster mitigation.
UNESCO’s work on Geohazard Risk Reduction operates in accordance with the four
Priorities for Action of the Sendai Framework for Disaster Risk Reduction 2015-2030:
Priority 1: Understanding disaster risk
Priority 2: Strengthening disaster risk governance to manage disaster risk
Priority 3: Investing in disaster risk reduction for resilience
Priority 4: Enhancing disaster preparedness for effective response and to Build
Back Better in recovery, rehabilitation and reconstruction.
Mitigation measures for risks associated with geohazards can broadly be classified in six
1. land use plans
2. enforcement of building codes and good construction practice
3. early warning systems
4. construction of physical protection barriers
5. network of escape routes and "safe" places
6. community preparedness and awareness building.
Early warning systems and construction of physical protection barriers have been singled
out as specific tasks in the proposed ICG research. Together with the other four categories,
they form the backdrop for a mitigation strategy.
The results of hazard and risk mapping and analyses will be used to formulate mitigation
strategies to assist decision-making on the need and cost-benefit of hazard mitigation
works. Based on such a strategy, protection measures can be developed and their
cost-effectiveness and environmental soundness compared.
Figure 1. Examples of tsunami risk mitigation master plan. Left: Patong City: elevated
green-belt areas approximately 400 m inland from the beach, serving as safe escape hills,
and system of escape routes; car traffic to be banned from escape routes. Right: Ban Nam
Khem fishing village: layout of protection dike around Ban Nam Khem with escape routes
to safe high areas (Karlsrud et al. 2006).
A mitigation strategy would involve:
1. identification of possible disaster triggering scenarios, and the associated hazard
2. analysis of possible consequences for the different scenarios;
3. assessment of possible measures to reduce and/or eliminate the potential
consequences of the danger;
4. recommendation of specific remedial measure and if relevant reconstruction and
rehabilitation plans;
5. transfer of knowledge and communication with authorities and society.
The strategy developed by ICG and NGI for the tsunami-affected areas of Thailand after
the 26th December 2004 Indian Ocean tsunami provides a good example of what can be
Any mitigation strategy needs to be adapted for different natural hazards and different
parts of the world. Especially for developing countries, it is vital to establish and promote
proper land-use planning and construction practices to regulate human activities that
increase risk to earthquakes, landslides or tsunamis and to prevent settlement of
communities in high-risk areas.
The research performed by ICG will be part of the strategy plans and the communication
programs. Ensuring that people do not live in "high risk" zones will be included in the
decision process. As for physical protection measures, an "how to" and "do's and don't's"
guideline will be prepared, as well as recommendation for "best practice".
DigitalGlobe’s Quickbird satellite captured an image of the devastation around Kalutara,
Sri Lanka (top), on December 26, 2004, at 10:20 a.m. local time—about an hour after the
first in the series of tsunami waves hit. Credit: DigitalGlobe
Chilean 1960 earthquake/tsunami is considered the largest – or most extreme – geohazard
and natural disaster in modern history. Since there were several warning foreshocks the
earthquake itself did not take that many lives, but the tsunami came as a surprise and in
turn led to the construction of the Pacific tsunami early warning system.
The 7.8 magnitude earthquake in Tangshan, China, in 1976, is the most deadly earthquake
ever recorded. The number of deaths is however unclear (I've seen between 250 000 - 800
000) to date as the Chinese for political reasons towards the end of the Cultural Revolution
did not want to deal with the disaster other than saying that the disaster stricken would and
should rescue themselves etc.
Really extreme geohazards – megastunamis and supervolcanoes
Norway is situated in a safe distance from the Ring of Fire. Crustal movements in this part
of the world are very slow stemming from post glacial rebound, the uplift of ground due to
the absence of heavy glaciers that melted thousands of years ago. Norway has in fact a
rather high number of earthquakes as well but these far from qualify as extreme
geohazards. But, if we look at the geological history of Norway we find evidence of a
really extreme geohazard. More than 8000 years ago, the submarine Storegga landslide
caused a wide ranging megatsunami hitting most of our entire coastline.
To develop reliable early warning systems, the physical processes
and mechanisms need to be understood and methods need to be
developed for measuring, modelling and predicting geohazards, for
example landslides and tsunamis.
Developing early warning systems also requires:
setting criteria for parameters to be monitored and threshold values;
developing monitoring equipment and systems;
coordinating satellite radar data with local monitoring stations;
planning monitoring programs for high-risk areas;
developing computer-aided decision-making tools with e.g. mobile
data mapping and retrieval, and information management using
geographical information technology (GIT), Remote sensing (RS)
and 3D modelling.
Figure 2. Block diagram of a typical early warning system (DiBiagio & Kjekstad 2007).
In particular, criteria will have to be established, for example, for the
rate and scale of ground movements in vulnerable locations, and links
will have to be established between ground movement, rainfall and
groundwater levels that can be used to develop a methodology for
landslide forecasting. An early warning system can also be used to
"measure" the effectiveness of landslide management strategies.
ICG aims to develop new techniques in remote sensing for detailed
investigation and monitoring of, for example, large rock-slope
instabilities and failures (lidar, radar, remote sensing) and slope
instability following a flood, including methods and tools for high
resolution digital elevation models (DEM) analysis.
Passive seismic monitoring techniques will be adapted to monitoring
of potentially unstable rock slope sites, providing options for an early
warning system and vital information for the general understanding of
rock-slope failure, and its dynamics. A passive monitoring system was
installed at Åknes in the fall of 2005. The data are being analysed in
real-time, and will be integrated with other continuous measurements
into an early warning system.
ICG will prepare user-guidelines for data review, alarm facility and
follow-up, telemetry links, and actions to be taken in the event of
threshold values being exceeded. Logical diagrams (flowcharts) for
the interpretation of the monitoring and early warning system will be
developed and tested before they are released for use.
Early warning systems are quite target-specific, depending on the
hazard type and the local conditions. For example, earthquake
prediction (in the strict sense) is not yet within reach, so for the
foreseeable future, developing a "warning system" for earthquakes is
not a realistic mitigation strategy. For tsunamis, however, the situation
is different and more promising, even if the short warning times are
still a major challenge. A few minute to one-hour tsunami warning
would have saved many lives in December 2004.
The dimension of both the Storegga slide
itself and the resulting tsunami is almost
incomprehensible. We cannot even begin to
think what damage such a tsunami would
do to Norwegian oil industry, fisheries and
our coastal population if it would have
taken place today.
Moving over the North Atlantic and almost
to the West coast of the US in Wyoming, we find the beautiful Yellowstone national
park. Yellowstone is known for its wildlife and its many geothermal features such as
geysirs. This park is namely situated on the top of a vast calderas from several
volcano eruptions so big that Yellowstone merits the name supervolcano. Yellowstone
is monitored by scientists that for obvious reasons find this place particularly
interesting, and they report that there are no signs that indicate the supervolcano is
about to erupt any time soon.
Buildings need to be designed (and placed in locations) to withstand the
impact forces of geohazards and to provide safe dwellings for people. Land
can also be elevated to ensure that buildings are above a critical height, for
example to protect against tsunami danger.
Physical protection barriers may be used to stop or delay the impact of the
geohazards, reduce the maximum reach of its impact, or dissipate the
energy of the geohazards. On land, such barriers may include "soft"
structures in the form of dikes or embankments, or "hard" structures like
vertical concrete or stone block wall. Offshore, the structures could be
jetties, moles or breakwaters, or even submerged embankments. Any
measures need to be part of a community's master plan and subjected to
analyses to assess and circumvent any negative environmental impact.
If a well functioning and efficient warning system is in place, warning and
escape are probably the best way to prevent loss of life due to geohazards.
Developing functional networks of escape routes and safe places could
include a number of different measures, strongly dependent on the local
Area, village or city analyses should provide maximum tolerable distance
from buildings and activities to a safe place, and assess how to achieve this
maximum distance. Distances between buildings and safe areas could be
shortened by reducing the escape routes, or by establishing new safe areas
as artificial escape hills and safe buildings that are accessible to people at
ICG will contribute to the development of templates for communities to
assess and select physical protection measures. The above descriptions
are only examples of possible measures. A multitude of considerations
need to be taken into account when preparing templates that are to be
implemented in real-life cases. Local conditions are determinant in many
cases. A "how to" and "do's and don't's" guideline will be prepared.
A recommendation for "best practice" for physical protection measures was
prepared towards the end of the second five-year period of ICG, in the
EU-sponsored FP7 Project, SafeLand.
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