Chapter 2: Earth Materials

Learning Objectives

The goals of this chapter are to:

  • Identify common minerals based on physical properties
  • Classify rocks based on texture and mineral composition
  • Interpret environments in which igneous, sedimentary, and metamorphic rocks are formed
  • Reconstruct the transport and depositional history of sedimentary rocks

2.1 Introduction

are the basic building blocks of rocks, which means rocks are made up of different combinations of minerals or just one mineral in some cases. Figure 2.1a is an example of a rock called , which is made up of a combination minerals. A mineral is a naturally occurring, usually inorganic, solid that can be defined by a chemical formula and a crystal structure. Figures 2.1b and 2.1c are two other types of rocks, sandstone and gneiss, that contain the same minerals as the granite in Figure 2.1a. Even though they contain the same minerals, all three of these rocks formed in different ways. We’ll get to that a little later, so for now, let’s focus on minerals.

Granite, arkose sandstone, and gneiss with potassium feldspar, quartz, plagioclase, and biotite.
Figure 2.1 – These are three different types of rocks that contain the same minerals: A) granite; B) arkose sandstone; C) gneiss. Pink minerals are K-feldspar, clear minerals are quartz, white minerals are plagioclase feldspar, and black minerals are biotite. Image Credit: James St. John, A, B, and C.

2.2 Minerals

According to the International Mineralogical Association, there are 5,575 known minerals, but most rocks are composed of just a few common minerals. In fact, only seven minerals make up 84% of Earth’s crust (Figure 2.2). So when you find an unknown rock, there is a good chance that at least one of those seven minerals will be present.

Pie chart of mineral distribution in Earth's continental crust.
Figure 2.2 – This pie chart shows the distribution of minerals in continental crust. Silicate minerals, structures based around the silicon-oxygen tetrahedron, are the most abundant, making up 92% of all minerals in Earth’s crust. Image Credit: Daniel Hauptvogel, CC BY-NC-SA.

Geologists can take rocks and minerals back to their laboratories and run in-depth chemical analyses to determine what minerals they have collected, but you don’t have that luxury when you’re working in the field or the classroom. Instead, many minerals can be identified based on a few observable properties that don’t require high-tech, expensive equipment. So, distinguishing these seven common minerals, and others, is easier than you think. The common, observable properties of minerals are:

  • Color
  • Luster
  • Streak
  • Hardness
  • Cleavage/Fracture


The color of minerals comes from small impurities in the chemical composition. Sometimes elements substitute for one another in the atomic structure of minerals, which can lead to significant changes in color. It is perhaps the most straightforward observable property, but it is rarely a diagnostic property for most minerals because the same mineral can come in a variety of different colors. Quartz, for example, can come in just about any color (Figure 2.3).

Colors of quartz including colorless, white, dark gray, purple, yellow and pink.
Figure 2.3 – Quartz in six color varieties. Color may not be a diagnostic property for mineral identification. Other colors include green, red, orange, and brown. Image credit: Karla Panchuck, CC BY-NC-SA.


The term describes how light interacts with the surface of a mineral. There are many different terms used to describe the luster of a mineral, but for this lab manual, we are only concerned with metallic and non-metallic lusters (does it look like a metal or does it not look like a metal, respectively). Be careful, though, because shiny doesn’t always mean it’s metallic. The technical difference between the metallic and non-metallic is that metallic minerals do not allow light to pass through the atomic structure, and non-metallic minerals do allow some light to pass through.

Metallic luster in gold-colored pyrite with observable cleavage.
Figure 2.4a – Pyrite, a mineral with a metallic luster. Image credit: Wikimedia Commons user JJ Harrison, CC BY-SA.
Non-metallic luster in prismatic quartz with six-sided euhedral crystal form.
Figure 2.4b – Quartz, a mineral with a non-metallic luster and six-sided crystal form. Image credit: Wikimedia Commons user Didier Descouens, CC BY-SA.


Streak is the color of a mineral in powdered form, which is observed by scratching the mineral across a ceramic streak plate (Figure 2.5). The powder color can be different than the mineral color because, when in powdered form, the mineral’s impurities do not have a significant effect on the absorption of light. The way light is absorbed and reflected is what gives things color. Streak is only a diagnostic property for metallic minerals because most other minerals have the same streak color as the hand-sample color, or the streak is white.

Streaks of hematite (red-brown), magnetite (black), sphalerite (honey-brown), and galena (dark gray) on a streak plate.
Figure 2.5 – Four different dark grey metallic minerals with varying streak colors. The minerals are from upper left clockwise: hematite (red-brown), magnetite (black), sphalerite (honey-brown), and galena (dark gray). Image credit: Karla Panchuck, CC BY-NC-SA.


The hardness of a mineral is a mineral’s ability to resist abrasion or scratching, and is determined using Mohs’ scale of mineral hardness (Figure 2.6). Mohs’ scale is qualitative, which means there is no quantitative relationship between hardness values; a 10 on the Mohs scale is not 10 times harder than a 1. The scale is based on the relative relationship between 10 different minerals.

How do you determine mineral hardness? When identifying mineral hardness in the classroom, you can use common objects that have a hardness value defined on the Mohs scale. These common objects and their hardness are located on the right side of Figure 2.6.

Mohs scale of relative hardness of minerals compared to common objects (finger nail, copper penny, glass, steel nail, and masonry bit).
Figure 2.6 – This is the Mohs scale of relative mineral hardness that uses 10 minerals with no quantitative relationship between the values. Lower numbers mean softer minerals; higher numbers mean harder minerals. On the right side is a list of common objects and their Mohs hardness that are used to test unknown minerals. Note: manufacturing differences can make the hardness of these common objects vary slightly. Image credit: Virginia Sisson, CC BY-NC-SA.


Cleavage is a term used to describe minerals that break in a predictable way. When minerals break along planes in which chemical bonds are weak, they produce flat surfaces called cleavage planes. Minerals can have 1, 2, 3, or more planes of cleavage (Figure 2.7). When there is more than one plane, the angle between the planes can also be described. Cleavage is one of the most challenging concepts for introductory geology students, but it is a useful property to identify minerals. For example, amphibole and pyroxene have the same physical properties except for cleavage. Knowing whether a rock contains amphibole or pyroxene could reveal a lot about the geologic history of that rock, such as where and how it most likely formed.

While the concept of mineral cleavage is difficult, identifying cleavage in minerals doesn’t have to be. When looking at a mineral, the easiest way to identify cleavage is to shine a light on its surface. If that surface brilliantly reflects light, then you most likely have a cleavage plane. Today, just about everyone has a cell phone with a flashlight that can be used to identify cleavage. In the absence of that, you can move the mineral around underneath an overhead light or use sunlight.

Minerals don’t always break in predictable ways, and that’s when we use the term fracture. Fracture describes irregular breakage in minerals, which usually occurs on mineral planes that don’t have any weak bonds. This sounds easy enough; flat surfaces mean cleavage and uneven surfaces mean fracture, but the truth is that it can be messy. Minerals can appear to have flat surfaces that are not actually cleavage, and fracturing can occur along planes of cleavage. This is why you need to look closely at the surface of a mineral using a hand lens.

A hand lens is a magnifying glass to help geologists look closely at rocks and minerals. To use the hand lens properly, you need to place the lens up against your face and then move the mineral or rock sample back and forth to get it into focus. A common mistake students make is holding the mineral or rock still and moving the hand lens; this is incorrect and is not going to help you see small-scale features in rocks and minerals.

Another thing that makes identifying cleavage difficult is the entire mineral sample does not have to take on the shape of the cleavages. The minerals in Figure 2.7 are examples where the samples do take on the cleavage shape, but for most minerals, this is not the case. Instead, you need to use a hand lens to look at the small details of a single side of the mineral. There will be slight changes in the heights of the surface that you wouldn’t be able to see without the hand lens. It is along those “elevation” changes that you should be looking for signs of cleavage. Then repeat your examination on the other sides.

Five common mineral cleavages showing the number and relative angle of cleavage planes, each with a representative minerals.
Figure 2.7 – Types of cleavage with mineral examples. This differ by the number and angle between cleavage planes. Image credit: cleavage models adapted from M.C. Rygel; muscovite from Daniel Hauptvogel, all other minerals from James St. John, CC BY-SA.

Exercise 2.1 – Identifying Common Minerals

Your instructor will provide you with a set of unknown minerals. Use either Table 2.1 as your guide or the mineral flow charts in Figures 2.8, 2.9, and 2.10, identify each mineral using its physical properties. Fill in this information in Table 2.2. There are also many online guides to help you identify minerals such as MineralID, Smart Geology, or the app Mineral Identifier.

Table 2.1 – Physical properties of common minerals
Hardness Color Cleavage Other Mineral Name
1-2 Colorless One plane “Chewy” Clays
2-2.5 Colorless, light tan, yellow One plane Can be peeled into transparent, flexible sheets Muscovite
2-2.5 Brown, black, green One plane Can be peeled into thin, flexible sheets Biotite
2.5 Colorless, white Three directions at 90° Cubic crystals, salty taste Halite
3 Colorless, white Three directions not at 90° Rhombic crystals, reacts with HCl, double refraction Calcite
5-6 Dark green to black Two directions at ~60° and ~120° Elongated crystals Hornblende (Amphibole)
5-6 Dark green to black Two directions at ~90° Elongated crystals Augite (Pyroxene)
6 Colorless, pink, gray Two directions at 90° Stubby, prismatic crystals Potassium Feldspar
6 Colorless, white, gray, black Two directions at 90° Can have small grooves on one of the cleavages called striations Plagioclase Feldspar
6.5-7 Green No cleavage, can have conchoidal fracture Stubby crystals or granular masses Olivine
7 Mainly red and black Conchoidal fracture May occur in 12-sided, circular crystals Garnet
7 Varied Conchoidal fracture May occur in six-sided, elongated crystals Quartz
Flowchart for light-colored minerals with nonmetallic luster
Figure 2.8 – Flowchart for light-colored minerals with nonmetallic luster. Image credit: Virginia Sisson, CC BY-NC-SA.
Flowchart for dark-colored minerals with nonmetallic luster
Figure 2.9 – Flowchart for dark-colored minerals with nonmetallic luster. Image credit: Virginia Sisson, CC BY-NC-SA.
Flowchart for minerals with metallic luster
Figure 2.10 – Flowchart for minerals with metallic luster. Image credit: Virginia Sisson, CC BY-NC-SA.
Table 2.2 – Worksheet for Exercise 2.1
Sample ID Color Hardness Cleavage Mineral Name

2.2 Rocks

Rocks are what most people think of when they hear “geology”. Rocks are made up of different minerals, and in some cases, just one mineral. Minerals in rocks can be difficult to identify, though. Even upper-level geology majors and practicing geoscientists can have trouble identifying minerals in rocks. Nonetheless, it is an essential skill to practice because it serves as one of the primary ways we can identify rocks. When geologists go into the field to study rocks, one of the first things they do is get up close and personal with the rock, usually with a hand lens, and identify what minerals are present. Let’s practice with a simple exercise on identifying some common minerals in rocks in Exercise 2.2.

Exercise 2.2 – Identifying Minerals in Rocks

Your instructor will provide you with a selection of different rock samples. From the list of minerals in Table 2.3 below, indicate which rock sample(s) contain those minerals by filling out the “Rock Sample” column.

Table 2.3 – Worksheet for Exercise 2.2
Mineral Rock Sample
Plagioclase feldspar

Based on how rocks are formed, geologists classify them into three basic types: igneous, sedimentary, and metamorphic. Igneous rocks form when maga or lava cools and solidifies. Minerals start to crystallize as the magma cools and they interlock with each other in random orientations. Sedimentary rocks form when broken pieces of other rocks from weathering processes get cemented together. They can also form when minerals precipitate out of a solution, typically water, and are cemented together through sedimentary processes. Metamorphic rocks are pre-existing rocks that are altered by heat and pressure. Metamorphism often results in some type of pattern or orientation to the minerals, called foliation.

Exercise 2.3- Identifying Types of Rocks

Your instructor will provide you with a selection of unknown rocks. Using what you know about how each rock type forms, determine the rock type for each of your unknown samples in Table 2.3.

Table 2.4 – Worksheet for Exercise 2.3
Rock Type Sample

2.3 Igneous Rocks

The classification of igneous rocks is based on chemistry and texture, which means it’s based on what minerals they have and how large the mineral grains are (Figures 2.11 and 2.12). The chemistry and temperature of the magma determine which minerals will form, and the size of the mineral grains is determined by how quickly the magma cools. Slow cooling rates (thousands to millions of years) make larger mineral grains, fast cooling rates (days to hundreds of years) produce smaller grains. Igneous rocks with large grains are called intrusive or plutonic rocks, which means the magma cooled slowly within the earth. Igneous rocks with tiny grains are called extrusive or volcanic because the lava cooled on the surface of the earth. The distinction between small and large grains is whether you (or a well-trained geologist) can identify the minerals without using any magnification. Using this information, if you had a coarse-grained igneous rock that contained 10% quartz, 10% potassium feldspar, 50% plagioclase feldspar, 20% pyroxene, and 10% amphibole, it would be called a diorite. If the mineral grains were fine (small), it would be called an andesite.

How would you identify minerals in an extrusive igneous rock if the grains are too small to see? There are a couple of tricks that geologists use that can help you identify which type of igneous rock you have. Do you remember how color was not a diagnostic property for minerals? Well, color in an igneous rock can be diagnostic! Iron and magnesium bearing minerals are darker in color, like olivine and pyroxene, and they are the primary minerals that make up mafic and ultramafic igneous rocks. Potassium, aluminum, and silica-rich minerals, like quartz and potassium feldspar, are lighter in color and make up intermediate and felsic igneous rocks. Geologists can use the color of igneous rocks, and the relative amount of light- vs. dark-colored minerals to help identify them (Figure 2.12).


Igneous rock compositions, their mineral content, and examples of extrusive and igneous rocks.
Figure 2.11 – Igneous rock classification table. To use this table, you will need to determine the relative abundance of seven common minerals found in igneous rocks. The bottom two rows show intrusive and extrusive igneous rocks. Image credit: Karla Panchuk after Steven Earle, CC BY-SA-NC.
Top is how to classify igneous rocks using percentages of light and dark colored minerals. Bottom shows relative proportions of light versus dark minerals.
Figure 2.12 – Identifying igneous rocks using mineral color. a) The relative percentages of light (felsic) and dark-colored minerals (iron and magnesium rich) minerals, b) Images you can use to estimate percentages of light- and dark-colored minerals. Note: our eyes tend to overestimate the amount of dark colors, so be careful of this when looking at the rocks. Also, for low abundances of accessory minerals, their size will influence your results. Image credit: a) Karla Panchuk after Steven Earle, CC BY; b) Virginia Sisson, CC BY-NC-SA.

Exercises 2.4 – Identifying Igneous Rocks

Your instructor will provide you with a set of unknown igneous rocks to identify. Use Figures 2.11 and 2.12 to fill out the information in Table 2.5 below.

Table 2.5 – Worksheet for Exercise 2.4
Sample Intrusive or Extrusive % Light vs. Dark Minerals (intrusive) or color (extrusive) Rock Name

2.4 Sedimentary Rocks

Sedimentary rocks can form in several ways (Figure 2.13) and are classified as clastic, chemical, and organic based on how they form. The most common way is when other rocks weather into small particles, and these particles are transported by wind, water, or ice to an area where they are deposited and layers accumulate. These are called clastic sedimentary rocks and are primarily classified based on their grain size. Sedimentary rocks can also form through chemical or organic processes. If you have a body of water that contains dissolved salt and then that water evaporates, it will leave behind the salt as a solid mineral form called a precipitate, which is a chemical sedimentary rock. Organic sedimentary rocks form from the accumulation of organic debris. This organic material would need to be buried very quickly so it doesn’t decay. Organic material in sedimentary rocks is where coal, oil, and natural gas come from. Then, there are sedimentary rocks that we can classify as bio-chemical. For example, foraminifera are tiny, single-celled organisms that build shells out of calcium carbonate. When they die, their shells can sink to the bottom of the ocean and accumulate to form a sedimentary rock. In this case, biological processes caused the chemical precipitation of the mineral for the shell, but it doesn’t contain organic compounds, so we call it bio-chemical.

In all cases, sediments need to be buried so the lithification process can take place. There are two steps involved in turning sediment into sedimentary rock; compaction and cementation. When grains are deposited, they typically have empty spaces between them called pores (porosity). Compaction reduces porosity by bringing the grains closer together. The best analogy to think about is your garbage can at home. When the garbage can is full, do you change the bag right away, or do you push down on it in order to create more room? By pushing down on trash, you reduce the space between the pieces of garbage and create more room at the top; this is compaction. Sediments work the same way, but what’s causing the compaction is the accumulation of more sediment on top of the previously deposited material. The more sediment that gets piled on top, the more compact the sediment beneath it becomes. Cementation is a process that happens during compaction. As sediments become compacted, water that is in the pore spaces gets squeezed out of the sediment. As the water leaves the sediment, it precipitates minerals into those pore spaces acting as a glue that holds the sediment together. The common minerals that make up the cement are calcite, quartz, and pyrite.

Figure 2.13 – NOT FINAL

The locations where sediments are deposited are called depositional environments. There are a variety of depositional environments (Figure 2.14) that can be broken up into three broad categories: continental, marine, and transitional. Continental environments are areas where sediments are deposited on continents, marine environments are areas where sediments are deposited in the ocean, and transitional environments are coastal and tidal locations between continental and marine environments. Table 2.6 contains the most common environments, a brief description of them, and the common sedimentary rocks that can form in those environments.

Block diagram of primary depositional environments of sediment.
Figure 2.14 – Common depositional environments for sediments ranging from continental to marine settings. Image credit: Modified from PePeEfe after Mikenorton, CC BY-SA.


Table 2.6 – Depositional Environments for Sedimentary Rocks
Environment Description Common Types of Sedimentary Rocks
Aeolian Sediment deposited by wind; primarily deserts and coastal regions; well-sorted sand; can be red in color; variable energy. Example in Google Earth: Algeria. Sandstone
Alluvial Fan-shaped deposits caused by moving water; usually found in arid or semi-arid regions; contains gravel, sand, silt, and/or clay; poorly sorted; high energy; creates alluvial fans. Example in Google Earth: Death Valley National Park, California. Conglomerate, breccia, sandstone, shale
Fluvial Sediment deposited by moving water, primarily rivers; can contain gravel, sand, silt, and/or clay depending on how fast the water moves; variable energy; commonly red in color from oxidation. Example in Google Earth: Upper Mississippi River in Illinois/Missouri. Conglomerate, sandstone, shale
Lacustrine Lake settings that can contain sand, silt, or clay; generally low energy. Example in Google Earth: Lake Winnipesaukee, New Hampshire. Sandstone, shale
Glacial Sediment deposited by glaciers; variable grain sizes; poorly sorted. Example in Google Earth: Southern Patagonia, Argentina. Conglomerate, breccia, sandstone, shale
Evaporitic Forms where water evaporates and leaves behind mineral precipitates. Example in Google Earth: Utah. Limestone, rock salt, rock gypsum.
Beach Along coastlines, sediment transported by wave action; contains well-sorted gravel and sand; high energy. Example in Google Earth: Island Beach State Park, New Jersey. Sandstone
Deltaic Where a river empties into a body of water; contains gravel, sand, and silt; low energy. Example in Google Earth: Yukon River, Alaska. Sandstone, shale
Lagoonal A shallow body of water separated from a larger body of water by barrier islands or reefs; very low energy; contains silt and clay. Example in Google Earth: East Matagorda Bay, Texas. Limestone, shale, coal (swamps)
Tidal Affected by the tides; mainly silt and clay; can create tidal flats; low energy. Example in Google Earth: Bay of Fundy, Nova Scotia. Shale
Shallow Located on the continental shelf; mainly sand and silt; energy decreases with distance from shore. Example in Google Earth: Eastern Gulf of Mexico. Sandstone, shale, limestone
Deep The deep ocean; very low energy; mainly clay; can contain turbidite deposits of variable sediment sizes. Example in Google Earth: Pacific Ocean. Shale, chert, limestone
Reef A bar of rock, sand, or coral. Example in Google Earth: Great Barrier Reef, Australia. Limestone

You’ll notice that the same type of sedimentary rock can be found in many different settings, which can make it difficult to determine the depositional environment. Sometimes, the color of a sedimentary rock can be an indicator of the depositional environment because the color is largely determined by how much oxygen is available as the sediment is buried and lithified (Figure 2.15). Reddish colors can indicate oxygen-rich continental environments, like a flowing river or desert. Green and light gray colors mean low oxygen, which can be found in shallow marine environments. Black color means no oxygen, which would indicate a deep marine or swamp environment. Color is not always a good indicator, though, because the sediment only alters its color after it is buried, and the conditions affecting buried sediment may not always match the environment on the surface above it.

Depositional environment of sedimentary rocks based on rock color from black to gray to green to red-brown.
Figure 2.15 – Typical colors of sedimentary rocks based on their depositional environment and oxygen availability during burial and lithification. Image credit: Daniel Hauptvogel, CC BY-NC-SA.

The size and shape of grains within sedimentary rocks can also help you interpret the history of that sediment (Figure 2.16). When all of the grains in the rock are about the same size, it is called well-sorted. Usually, sediment needs to travel a far distance from its source to be well-sorted. In contrast, poorly sorted sediment is very close to its source. The rounding of sediment grains is also an indicator of how far the sediment traveled. Grains that have smooth edges are considered well-rounded and have traveled a far distance. These grains start out with rough edges and become smoother as they travel further and further. In contrast, grains with sharp edges are considered angular and have not traveled a far distance.

Characterizing the sorting and roundness of sediment grains.
Figure 2.16 – Sediment grain characteristics and the relative distance traveled from their source rock. a) Sorting of sediment; b) Sediment roundness often called sphericity. Image credit: a) Adapted from Wikimedia user Woudloper, CC BY-SA; b) Adapted from Wikimedia user Woudloper, CC BY-SA.

Exercises 2.5 – Identifying Sedimentary Rocks

Your instructor will provide you with a set of unknown sedimentary rocks to identify. Use Figures 2.13 to 2.17 and Table 2.2 to fill out the information in Table 2.7 below.

Table 2.7 – Worksheet for Exercise 2.5
Sample Clastic, (Bio-) Chemical, or Organic Rock Name Possible Depositional Environments
Figure 2.17 – Flowchart for identifying sedimentary rocks. Image credit: Virginia Sisson, CC BY-NC-SA.

2.5 Metamorphic Rocks

Metamorphic rocks form when any pre-existing rock is altered by intense heat and/or pressure. The two sources of heat for metamorphism are the heat from a magma chamber and the geothermal gradient, which is the natural increase in temperature when getting deeper into the earth. Pressure also increases with depth in the earth, but intense pressure also occurs at tectonic plate collisions. This is why most of the world’s metamorphic rocks are found in mountain belts or ancient mountain belts.

Heat and pressure can cause a number of changes in rocks, such as recrystallizing minerals, creating new minerals, and orientating minerals in a direction perpendicular to pressure. The orientation of minerals is the most noticeable feature when recognizing metamorphic rocks. Heat works to “soften” minerals, which can allow ions to begin migrating in and out of crystal structures, and then pressure forces the minerals to re-orientate. Heat is the most important factor, though; without heat, the rock would just break under the intense pressure. When minerals take on some type of orientation due to metamorphism, it’s called foliation. As pressure increases, so does the degree of foliation. Not all metamorphic rocks exhibit foliation because some minerals change very little under metamorphic conditions (quartz, for example); these types of metamorphic rocks are called non-foliated and can be very difficult to tell apart from igneous rocks.

The pre-existing rock, before metamorphism occurred, is called the protolith. If you’ve taken an introductory Physical Geology class, you learned to differentiate metamorphic rocks mostly by their texture and classified them as slate, phyllite, schist, and gneiss. You may have even been introduced to the term migmatite for rocks that had both igneous and metamorphic textures as these were starting to undergo partial melting. These textural terms are associated with changes in metamorphic temperature and pressure, often referred to as the metamorphic grade or just grade.

Figure 2.18 – Metamorphic rock classification by texture, protolith, mineralogy, and type of metamorphism. Also shown are the US Geological Survey map patterns for each rock type.  There are not standard patterns for either amphibolite or eclogite; so we made a pattern using the first letter of these rock names.  Also shown are photos of samples from the Wards Classic American rock collection except for gneiss, migmatite, hornfels, and eclogite which are all photos from Sisson (unpublished). Image credit: Virginia Sisson, CC BY-NC-SA.

Exercises 2.6 – Identifying Metamorphic Rocks

Your instructor will provide you with a set of unknown metamorphic rocks to identify. Use Figures 2.18 and 2.19 to fill out the information in Table 2.8.

Table 2.8 – Worksheet for Exercise 2.6
Sample Foliated or Non-foliated Rock Name Protolith
Figure 2.19 – Flowchart for identifying metamorphic rocks. Begin by deciding if the metamorphic rock exhibits any type of foliation. Then determine their grain size.  Next, look at other features such as mineralogy, hardness, and how the rock breaks. Image credit: Virginia Sisson, CC BY-NC-SA.

Let’s build on your metamorphic rock foundation and start to interpret metamorphic rocks using mineral assemblages. To do this, you have to start by knowing the protolith (parent) material for the metamorphic rock. If you can distinguish mafic (basaltic protoliths) from shale (sedimentary protoliths), then this will be easy.

During metamorphism, rocks that have a shale protolith will always have quartz, feldspar, and muscovite mica. So, you are on the lookout for new minerals. Most of these are not taught in Physical Geology, so here are some minerals you will need to know: chlorite (a dark green mica), biotite, cordierite, garnet, staurolite, sillimanite, kyanite, and andalusite (Figure 2.20). The last three minerals on this list you may remember are polymorphs: minerals with the same composition but unique internal structures (Figure 2.21). If you can identify these alumino-silicate minerals in a schist or gneiss, then you can determine the pressure-temperature conditions for that terrane. In general, rocks with kyanite and sillimanite are found in areas that have undergone continent-continent collision and are considered medium-pressure terranes. In contrast, areas with andalusite and sillimanite have an elevated geothermal gradient in response to either divergent zones or unusual ocean-continent collision zones and are called low-pressure terranes.

Index minerals for metamorphic rocks and their temperature ranges.
Figure 2.20 – Index minerals for metamorphic rocks and their temperature ranges. These index minerals are for metamorphosed shales (pelites). These rocks all have quartz, feldspar, and muscovite mica in addition to other index minerals. Green is for low-grade rocks that have chlorite. As temperature increases, biotite (brown) begins to form. As the temperature continues to increase, medium grade rocks will have garnet (red) and kyanite (blue). Both of these are easy to distinguish as these minerals grow as porphyroblasts and often form bumps in the rocks. Finally, at the highest grade, sillimanite appears, and then the rocks begin to melt, forming migmatites. Image credit: Virginia Sisson, CC BY-NC-SA.
Pressure temperature diagram for three polymorphs of aluminosilicate minerals (kyanite, andalusite, and sillimanite).
Figure 2.21 – Pressure temperature diagram for aluminosilicate minerals showing the stability fields for andalusite, kyanite, and sillimanite. The dashed line is defined by the stability of pyrophyllite and limits the lower temperature of andalusite and kyanite. This diagram was calculated using the data of Berman (1991). Image credit: Virginia Sisson, CC BY-NC-SA.

During a continental collision, the sequence of metamorphic rocks that forms can be predictable by following the Barrovian sequence (Figure 2.22). Basically, the further away from the collision or the higher up in the crust the rock is, the lower the grade of metamorphism while the closer and deeper the rocks are, the higher the grade. So, how do these deeper, high-grade metamorphic rocks become exposed at the surface? The answer is erosion; the low-grade rocks above need to be removed in order to expose the high-grade rocks below. This is relatively easy for an uplifting mountain belt because the increased surface area of the mountain allows more weathering and erosion to take place. If you are looking for exposures of high-grade metamorphic rocks, your best chance of finding them is near the center of the collision where much of the uplift and erosion took place. The Barrovian sequence pattern is when the grade of metamorphic rocks increases as you approach the collision center, reaching the highest grade near the center, and then decreasing in grade as you move to the other side (like a mirror image). Of course, this perfect, mirror-image pattern is not always seen because the rocks may not be exposed, or maybe faults have moved rocks from their original positions, or there was significant variability in pressure-temperature conditions.

Zones of metamorphic index minerals in a continental convergence tectonic setting.
Figure 2.22 – Metamorphic zones in continental convergence. The colored areas are the metamorphic zones; these use the same scheme as in Figure 2.20. Green is a chlorite, brown is for biotite, red is for garnet, blue is for kyanite, and gray is for sillimanite. This scheme does not include andalusite as a separate zone; it would overlap the kyanite zone. The dashed line is the position of the boundary between the continental crust and mantle, called the Moho. It is shortened from the term Mohorovičić discontinuity after the Croatian seismologist Andrija Mohorovičić (1857-1936) who discovered this. This figure was simplified from numerical models by Lyubetskaya and Ague (2010). The position of the Moho shows the initial temperature gradient, and the metamorphic zones reflect the temperature 30 million years after continental collision. Image credit: Virginia Sisson, CC BY-NC-SA.

Exercise 2.7 – Understanding the Barrovian Sequence.

Now let’s put together information from the rocks you’ve looked at for this exercise with a map of the metamorphism of the Scottish Highlands. British geologist George Barrow mapped the area in Figure 2.23 in the 1890s and realized that he could distinguish metamorphic zones by using their mineral assemblages.

Barrovian metamorphic zones of northern Scotland.
Figure 2.23 – Barrovian metamorphic zones in northern Scotland. The insert map of Great Britain (Ireland, Scotland, Wales, and Britain) shows the location of this map on the northeastern side of Scotland. The colors used for the different metamorphic zones are the same as in figure 2.20. Green is a chlorite, brown is for biotite, red is for garnet, blue is for kyanite and gray is for sillimanite. Barrow also mapped a staurolite zone which is not shown here. The black circles indicate the location of major Scottish cities. The area is bound by two faults, the Great Glen strike-slip fault to the northwest and the Highland Boundary reverse fault to the southeast. Image credit: Virginia Sisson, CC BY-NC-SA; simplified from Wikimedia user Woudloper, CC BY-SA.
  1. Label the colored areas on the map according to the scheme used in Figure 2.20.
  2. Where on this map do you think your samples from Exercise 2.6 occur?
  3. On the map, mark which areas you think were the shallowest and deepest during metamorphism.
  4. Draw a line on the map that corresponds to only a change in temperature.
  5. Using Figure 2.23, plot an approximate path through the metamorphic zones. This is called a pressure-temperature (PT) path. Geologists now like to add geochronology constraints (t) or compositional constraints (x), so you’ll hear them discuss PTt paths when trying to determine the tectonic history of a region.
  6. To help understand the metamorphic history of a region, geologists will construct geodynamic models of how heat and pressure in the earth evolve through time. These models start with an initial set of conditions for the area. In Figure 2.22, the heavy black lines show two continental plates colliding; each has an initial lithosphere thickness of ~ 35 km. The convergence causes deflection of the Moho (Mohorovičić) discontinuity between the crust and mantle. Over time, the temperature readjusts. The colored areas on this plot show the metamorphic assemblages after ~35 million years since collision (Lyubetskaya and Ague, 2010). The first step to using this diagram is to convert metamorphic pressures to depth. A good rule of thumb is that 100 MPa is ~3 km depth. Use your results from 2.2e to fill in Table 2.9.
    Table 2.9 – Worksheet for Exercise 2.7f
    Metamorphic Zone Pressure Depth Temperature
  7. Using your calculated depths from Table 2.9, draw a line on Figure 2.22 that corresponds to your sequence of metamorphic assemblages from the map.

2.6 Using Rock to Interpret Earth’s History

Igneous, sedimentary, and metamorphic rocks can all be used to interpret Earth’s history. Igneous rocks can help determine the plate tectonic setting of an area, sedimentary rocks help determine the environment of an area, and metamorphic rocks provide information about rocks prior to their deformation as well as the tectonic setting.

Exercises 2.8 – Interpreting Volcanic History

There are a few places on Earth with significant mountainous relief as well as igneous activity. One such place is the western margin of South America, which has been the site of ocean-continent convergence since the Jurassic (~185 Ma).  Thus, it has a long history of almost continuous subduction-related volcanism and igneous activity. In the Central Andes Mountains, there are igneous rocks that formed during the Neogene portion of geologic time between 23 and 2.5 million years ago; these include both intrusive and extrusive rocks (Figure 2.24). These volcanic rocks are overlain by more recent volcanoes.

This map shows igneous rocks (extrusive in orange, intrusive in red) and volcanoes from a region in Peru.
Figure 2.24 – Map of the Neogene igneous rocks of the Peruvian Andes. Intrusive rocks are shown as red areas and extrusive volcanic rocks are shown in the orange areas. Black triangles of locations of active volcanoes. Note that this map is not oriented north-south. Image credit: Virginia Sisson, CC BY-NC-SA; simplified from Pfiffner and Gonzalez, 2013.
  1. In the suite of igneous rocks provided by your instructor, which rocks are associated with the Neogene intrusives?
  2. Which rocks would be associated with the Neogene extrusive volcanic rocks?
  3. Would the same rock types be in the active volcanos?
  4. Describe at least three differences between the map patterns of the recent volcanos, Neogene extrusives, and Neogene intrusives.
  5. What can explain some of these differences in the map patterns?

Exercise 2.9 – Rocks of the Grand Canyon

The Grand-Canyon has amazed tourists and geoscientists alike, not only for its beauty, but by the fantastic stratigraphic sections (layers of sedimentary rock) exposed in its walls. The South Rim of the Grand Canyon is composed of metamorphic, igneous, and various sedimentary rocks. Figure 2.25 shows a stratigraphic column of the Grand-Canyon (see Figure 0.3 for a sketch of the Grand Canyon).

Rocks of the Grand Canyon, including names and symbols for the rock formations.
Figure 2.25 – Cross-section of the Grand Canyon. This view shows the inner gorge, Temple Butte, and southern rim on the west bank of the south-flowing Colorado River. The elevation of Temple Butte is 1618 m; the maximum depth of the Grand Canyon is over a mile or 1.829 km.. This cross-section shows simplified stratigraphy of the sedimentary, igneous and metamorphic rocks exposed in the cliffs and walls of the canyon. The different formations are shown using patterns for the different rock types that are commonly used by the U.S. Geological Survey. In other regions of the Grand Canyon, there is also the Grand Canyon Supergroup that overlies the Vishnu Schist and Zoraster granite. This is not shown in this figure. Figure is adapted from several publicly available stratigraphic columns including the National Park Service. Image credit: Virginia Sisson, CC BY-NC-SA.

The table below contains descriptions of some of the different rocks found in the Grand Canyon.

  1. Which rocks from your kit resemble the sedimentary units marked with a star on the Grand Canyon stratigraphic section? Fill this out in the third column Table 2.10.
  2. The symbols used to describe these rocks are standardized. Based on your findings from part a, complete the legend in Figure 2.25 by identifying the type of rock each symbol stands for.
Table 2.10 – Rock formations and descriptions for the Grand Canyon
Rock Formation* Description Matching Rock From Your Kit
Vishnu Group Group of foliated metamorphic rocks that are the result of various rocks, mostly igneous, being exposed to high degree metamorphism.
Zoroaster Group Igneous rocks composed of coarse-grained intrusions composed mostly of feldspar and quartz, with minor hornblende and biotite.
Tapeats Formation Composed mostly of sand-sized quartz grains with medium roundness. It is mostly massive, but bedding is visible in some areas.
Bright Angel Formation Composed of fine-grained shale, and trilobite fossils are common within this unit.
Mauv Formation Chemical and biochemical sedimentary rocks that react with dilute acid.
Hermit Formation Extremely fine-grained, dark-colored sedimentary rocks which have easily observable bedding. Most of the minerals are clay-minerals, but can’t be observed, even with the help of a hand-lens, due to their extremely fine grain size.

*See this Wikipedia entry for the differences in stratigraphic unit terminology.

Most of the rocks formed at divergent plate boundaries are igneous with a thin layer of deep marine sediments on top. Since these sediments accumulate deep below the ocean surface, there are only chert and marine clays. Instead, there are significant accumulations of igneous rocks derived from the decompression melting of Earth’s peridotite mantle. Do you recall what kind of rock forms from the partial melting of an ultramafic rock? It will be a mafic rock, either extrusive or intrusive. The oceanic crust is subdivided into five layers, as shown in Figure 2.26. This entire sequence is called an ophiolite. Ophiolites are typically exposed on Earth’s surface during the end of the subduction process as two continents move closer together and collide. During these final stages of subduction, the oceanic lithosphere gets stuck in the middle of the continental collision and is uplifted with the newly forming mountain belt. This is why most ophiolites are found in mountains.

Exercise 2.10 – Ophiolites

Ophiolites offer clues to where oceanic crust used to exist, and they help geologists understand processes at mid-ocean ridges.

  1. Your instructor has given you a suite of rocks that represent different layers in an ophiolite sequence. Identify your rocks and complete the legend at the bottom of Figure 2.26 by filling in the compositions and rock names.
    A sequence of igneous and sedimentary rocks found in ophiolite complexes.
    Figure 2.26 – Sequence of rocks in an ophiolite complex. Since it is difficult find a location with all of these igneous rock types exposed in one outcrop, this is a composite of several different ophiolite exposures. The patterns use for the different units are adapted from the U.S. Geological Survey. Since most ophiolite sequences are composed of primarily mafic to ultramafic igneous rocks that are underneath deep marine sediment, we used dark green to brown colors for the different rock types. Image credit: Virginia Sisson, CC BY-NC-SA.
  2. Explain why ophiolites show a sequence that moves from intrusive rock to extrusive.
  3. What is the significance of the pillow lavas compared to the sheeted dike layer?
  4. Now use your knowledge of ophiolites to answer some questions about the geologic history of the Himalayan mountains. On Figure 2.27, label the two tectonic plates. Are these oceanic or continental tectonic plates? What is the thickness of their lithosphere?
  5. What are the rocks shown in the black blobs on this geologic map? ____________________
  6. There are five other groups of rocks in the Himalayan Mountains. Each of these groups has a distinct geologic history.
      1. Why do you think these groups of rocks are distributed as long skinny units?
      2. The alignment of the rocks is called their strike direction. What is the strike direction of these groups? ____________________
      3. Does the strike change from east to west along the Himalayan Mountains? ____________________
      4. What can you infer from your observations?
  7. Make some observations about the map patterns of the ophiolites.  Are the ophiolite blobs in a continuous or discontinuous belt?
  8. Since ophiolites are signatures of previous ocean basins, what happened to all of the oceanic crust that was once between these two continental tectonic plates?
  9. Some geoscientists suggest that we use the term suture zone to describe where these occur during continent-continent convergence. Looking at the map, do you think this term is appropriate for the Himalayan Mountains. Explain your answer.
Geologic map of the Himalaya Mountain Range with a topographic insert.
Figure 2.27 – Geologic map of the Himalayan Mountain range. The insert is a shaded relief map of Asia including Russia and India. Most of the high topography is in an area called the Tibetan plateau. To the south is the low area of India and to the north is Russia and China. Map scale reference is at 30 degrees north latitude. Within the insert, is a black box that outlines the area shown below. The location for Mt. Everest (Chomolungma in Tibetan) is shown for reference. The geology map has six main lithologies. The northernmost unit includes igneous rocks of the Transhimalayan batholith and Kohistan arc shown in red stipple pattern. Scattered throughout the map are occurrences of ophiolite sequences shown as black areas. If you are at the eastern (right) edge of the mountains, the next unit is the Indus suture zone shown in a blue pattern. Next to this is the Tethyan zone in a brown stippled pattern. The next unit is Greater Himalayan sequence in wavy blue lines. The southernmost unit is the Lesser Himalaya shown as red dots. Each unit is separated from the next by either a thrust fault (line with barbed teeth) or normal fault (line with hatchures). There is one major strike-slip fault (Karakoram) that offsets part of the Himalayan mountains on the northwest side of the range. If you look carefully on the shaded relief map, you can pick out how far this unit goes into the Tibetan plateau. In this map are several letter abbreviations for NP (Nanga Parbat), A (Annapurna), K (Katmandu), E (Everest) and B (Bhutan). This map was simplified from MP Searle and PJ Treloar, 2019.

Exercise 2.11 – Southern Alaska

In southern Alaska, there is a sequence of metamorphic rocks called the Chugach Metamorphic Complex (Figure 2.28). In one area along the Tana River, you can walk upstream from phyllite to gneiss. However, to determine the metamorphic conditions, you need to identify the mineral assemblages. A team of geologists worked along this river for three summers and mapped the distribution of metamorphic minerals as well as the deformation, textural and geochronological history. The outcrops occur where a river along a glacier eroded away the glacial sediments and exposed the bedrock. The protoliths for these rocks were graywacke deposited in an accretionary prism during the Late Cretaceous. Now that you have the geologic history for this region, try to interpret the metamorphic pattern shown in Figure 2.28

This maps shows the metamorphic zones and textures of southern Alaska.
Figure 2.28 – Map of metamorphic zones in southern Alaska. The color scheme is similar to Figure 2.20 except that this area has andalusite instead of kyanite. Green is a chlorite, brown is for biotite, red is for garnet, blue is for andalusite, yellow for staurolite (a yellow-brown silicate) and gray is for sillimanite. The boundaries between these zones are isograds shown as solid lines. There are two dashed lines indicating boundaries between rocks with different textures (phyllite, schist and gneiss). The blue stippled pattern indicates the position of the Tana Glacier in 1988. The gray stipple pattern represents glacial deposits. The inset map shows an arrow pointing to the location of the Tana River in southern Alaska. This map is from adapted from Bowman et al. (2003) with unpublished mapping of Sisson.
  1. Do the textural boundaries between phyllite, schist, and gneiss correspond to mineralogical changes?  If not, why do you suppose they are different?
  2. Is this a medium- or low-pressure terrane? Did it form as a contact aureole or regional metamorphism?
  3. Using a graph of metamorphic temperatures, what is the temperature range from north to south along the Tana River?
  4. What is the horizontal geothermal gradient along the Tana River? To determine the horizontal geothermal gradient, determine the distance between the two points you used for estimating the metamorphic temperature.
  5. How does the horizontal geothermal gradient compare with vertical geothermal gradients in stable continental crust?
  6. What was the heat source for the metamorphism?
  7. This map shows the distribution of another mineral common in metamorphosed shales, staurolite. Using the temperature scale, determine the temperature that the staurolite formed.
  8. Using the pressure-temperature plot (Figure 2.21) which has the stability for the alumino-silicate polymorph minerals, plot the path that the highest grade rocks would have taken through time.

Exercise Contributions

Daniel Hauptvogel, Virginia Sisson, Carlos Andrade


Berman, R.G., 1991, Thermobarometry using multi-equilibrium calculations; a new technique, with petrological applications, The Canadian Mineralogist, v. 29 (4), p.  833–855

Bowman, J.R., Sisson, V.B., Pavlis, T.L., and Valley, J.W., 2003, Oxygen isotope constraints on fluid infiltration associated with high temperature–low pressure metamorphism (Chugach Metamorphic Complex) within the Eocene southern Alaska fore-arc, in Sisson, V. B., Roeske, S. M., Pavlis, T. P., editors, Geological consequences of ridge-trench interactions in the northern Pacific, Geological Society of America Special Publication no. 371, p. 237-252

Lyubetskaya, T., and Ague, J.J., 2010, Modeling metamorphism in collisional orogens intruded by magmas: I. Thermal evolution: American Journal of Science, v. 310, p. 427-458, DOI: 10.2475/06.2010.02

Pfiffner, O. A., and Gonzalez, L., 2013, Mesozoic–Cenozoic Evolution of the Western Margin of South America: Case Study of the Peruvian Andes, Geosciences 3, no. 2: 262-310, DOI: 10.3390/geosciences3020262

Searle, M.P., and Treloar, P.J., 2019, Introduction to Himamalyan tectonics: a modern synthesis, in Himalayan Tectonics: A modern synthesis ed by P.J. Treloar and M.P. Searle, Geological Society of London Special Publications, v. 483, p. 1-17


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