Lecture notes,2nd semester,unit-4

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1 Lecture notes,2nd semester,unit-4 VOLCANOES AND V OLCANIC ERRUPTIONS Since volcanic eruptions are caused by magma (a mixture of liquid rock, crystals, and dissolved gas) expelled onto the Earth's surface, we'll first review the characteristics of magma that we covered previously. Three basic types of magma: 1. Mafic or Basaltic-- SiO wt%, high in Fe, Mg, Ca, low in K, Na 2. Intermediate or Andesitic-- SiO wt%, intermediate. in Fe, Mg, Ca, Na, K 3. Felsic or Rhyolitic-- SiO %, low in Fe, Mg, Ca, high in K, Na. Gases - At depth in the Earth nearly all magmas contain gas. Gas gives magmas their explosive character, because the gas expands as pressure is reduced. Mostly H2O with some CO2 Minor amounts of Sulfur, Cl, and F Felsic magmas usually have higher gas contents than mafic magmas. Temperature of Magmas Mafic/Basaltic oC Intermediate/Andesitic oC Felsic/Rhyolitic oC. Viscosity of Magmas Viscosity is the resistance to flow (opposite of fluidity). Depends on composition, temperature, & gas content. Higher SiO2 content magmas have higher viscosity than lower SiO2 content magmas Lower Temperature magmas have higher viscosity than higher temperature magmas. Thus, basaltic magmas tend to be fairly fluid (low viscosity), but their viscosity is still 10,000 to 100,0000 times more viscous than water. Rhyolitic magmas tend to have even higher viscosity, ranging between 1 million and 100 million times more viscous than water. (Note that solids, even though they appear solid have a viscosity, but it very high, measured as trillions time the viscosity of water). Viscosity is an important property in determining the eruptive behavior of magmas. Magma Solidified Solidified Summary Table Chemical Temperature Viscosity Gas Content

2 Type Mafic or Basaltic Volcanic Rock Basalt Intermediate Andesite or Andesitic Felsic or Rhyolitic Rhyolite Plutonic Rock Composition Gabbro SiO2 %, high in Fe, Mg, oc Low Ca, low in K, Na Diorite SiO2 %, intermediate in oc Fe, Mg, Ca, Na, K Intermediate Intermediate Granite SiO2 %, low in Fe, Mg, Ca, high in K, Na High oc Low High The Products of Volcanic Eruptions Lava Flows When magma reaches the surface of the earth, it is called lava. Since it its a liquid, it flows downhill in response to gravity as a lava flows. Different magma types behave differently as lava flows, depending on their temperature, viscosity, and gas content. Pahoehoe Flows - Basaltic lava flows with low viscosity start to cool when exposed to the low temperature of the atmosphere. This causes a surface skin to form, although it is still very hot and behaves in a plastic fashion, capable of deformation. Such lava flows that initially have a smooth surface are called pahoehoe flows. Initially the surface skin is smooth, but often inflates with molten lava and expands to form pahoehoe toes or rolls to form ropey pahoehoe. (See figure 9.3d in your text). Pahoehoe flows tend to be thin and, because of their low viscosity travel long distances from the vent. A'A' Flows - Higher viscosity basaltic and andesitic lavas also initially develop a smooth surface skin, but this is quickly broken up by flow of the molten lava within and by gases that continue to escape from the lava. This creates a rough, clinkery surface that is characteristic of an A'A' flow (see figure 9.3e in your text). Lava Tubes - Once the surface skin becomes solid, the lava can continue to flow beneath the surface in lava tubes. The surface skin insulates the hot liquid lava form further cooling. When the eruption ends, liquid lava often drains leaving an open cave (see figure 9.3 in your text). Pillow Lavas - When lava erupts on the sea floor or other body of water, the surface skin forms rapidly, and, like with pahoehoe toes inflates with molten lava. Eventually

3 these inflated balloons of magma drop off and stack up like a pile of pillows and are called pillow lavas. Ancient pillow lavas are readily recognizable because of their shape, their glassy margins and radial fractures that formed during cooling (see figure 9.4b in your text). Columnar Jointing - When thick basaltic or andesitic lavas cool, they contract. The contraction results in fractures and often times results in a type of jointing called columnar jointing. The columns are usually hexagonal in shape. This often happens when lavas pool in depressions or deep canyons (see figure 9.4a in your text). Siliceous Lava Flows - High viscosity andesitic and rhyolitic lava flows, because they can t flow very easily, form thick stubby flows that don t move far from the vent. Lava Domes or Volcanic Domes - result from the extrusion of highly viscous, gas poor andesitic and rhyolitic lava. Since the viscosity is so high, the lava does not flow away from the vent, but instead piles up over the vent. Blocks of nearly solid lava break off the outer surface of the dome and roll down its flanks to form a breccia around the margins of domes. The surface of volcanic domes are generally very rough, with numerous spines that have been pushed up by the magma from below. Pyroclastic Material If the magma has high gas content and high viscosity, the gas will expand in an explosive fashion and break the liquid into clots that fly through the air and cool along their path through the atmosphere. Alternatively it blast out solid pieces of rock that once formed the volcanic edifice. All of these fragments are referred to as Pyroclasts = hot, broken fragments. Loose assemblages of pyroclasts called tephra. Depending on size, tephra can be classified as bombs. blocks, lapilli, or ash. Tephra and Pyroclastic Rocks Average Particle Size (mm) Unconsolidated Material (Tephra) Pyroclastic Rock >64 Bombs or Blocks Agglomerate 2 64 Lapilli Lapilli Tuff

4 <2 Ash Ash Tuff Blocks are angular fragments that were solid when ejected. Bombs have an aerodynamic shape indicating they were liquid when ejected. Bombs and lapilli that consist mostly of gas bubbles (vesicles) result in a low density highly vesicular rock fragment called pumice. Rock formed by accumulation and cementation of tephra called a pyroclastic rock or tuff. Welding, compaction and deposition of other grains cause tephra (loose material) to be converted into pyroclastic rock. Volcanic Landforms Volcanic landforms are controlled by the geological processes that form them and act on them after they have formed. Thus, a given volcanic landform will be characteristic of the types of material it is made of, which in turn depends on the prior eruptive behavior of the volcano. Here we discuss the major volcanic landforms and how they are formed Most of this material will be discussed with reference to slides shown in class that illustrate the essential features of each volcanic landform. Shield Volcanoes A shield volcano is characterized by gentle upper slopes (about 5o) and somewhat steeper lower slopes (about 10o). Shield volcanoes are composed almost entirely of relatively thin lava flows built up over a central vent. Most shields were formed by low viscosity basaltic magma that flows easily down slope away from the summit vent. The low viscosity of the magma allows the lava to travel down slope on a gentle slope, but as it cools and its viscosity increases, its thickness builds up on the lower slopes giving a somewhat steeper lower slope. Most shield volcanoes have a roughly circular or oval shape in map view. Very little pyroclastic material is found within a shield volcano, except near the eruptive vents, where small amounts of pyroclastic material accumulate as a result of fire fountaining events.

5 Stratovolcanoes (also called Composite Volcanoes) Have steeper slopes than shield volcanoes, with slopes of 6 to 10o low on the flanks to 30o near the top. The steep slope near the summit is due partly to thick, short viscous lava flows that do not travel far down slope from the vent. The gentler slopes near the base are due to accumulations of material eroded from the volcano and to the accumulation of pyroclastic material. Stratovolcanoes show inter-layering of lava flows and pyroclastic material, which is why they are sometimes called composite volcanoes. Pyroclastic material can make up over 50% of the volume of a stratovolcano. Lavas and pyroclastics are usually andesitic to rhyolitic in composition. Due to the higher viscosity of magmas erupted from these volcanoes, they are usually more explosive than shield volcanoes. Stratovolcanoes sometimes have a crater at the summit that is formed by explosive ejection of material from a central vent. Sometimes the craters have been filled in by lava flows or lava domes, sometimes they are filled with glacial ice, and less commonly they are filled with water. Long periods of repose (times of inactivity) lasting for hundreds to thousands of years, make this type of volcano particularly dangerous, since many times they have shown no historic activity, and people are reluctant to heed warnings about possible eruptions.

6 Cinder Cones Cinder cones are small volume cones consisting predominantly of ash and scoria that result from mildly explosive eruptions. They usually consist of basaltic to andesitic material. They are actually fall deposits that are built surrounding the eruptive vent. Slopes of the cones are controlled by the angle of repose (angle of stable slope for loose unconsolidated material) and are usually between about 25 and 35o. They show an internal layered structure due to varying intensities of the explosions that deposit different sizes of pyroclastics. On young cones, a depression at the top of the cone, called a crater, is evident, and represents the area above the vent from which material was explosively ejected. Craters are usually eroded away on older cones. If lava flows are emitted from tephra cones, they are usually emitted from vents on the flank or near the base of the cone during the later stages of eruption. Cinder and tephra cones usually occur around summit vents and flank vents of stratovolcanoes. An excellent example of cinder cone is Parícutin Volcano in Mexico. This volcano was born in a farmers corn field in 1943 and erupted for the next 9 years. Lava flows erupted from the base of the cone eventually covered two towns. Craters and Calderas Craters are circular depressions, usually less than 1 km in diameter, that form as a result of explosions that emit gases and ash. Calderas are much larger depressions, circular to elliptical in shape, with diameters ranging from 1 km to 50 km. Calderas form as a result of collapse of a volcanic structure. The collapse results from evacuation of the underlying magma chamber.

7 Crater Lake Caldera in southern Oregon is an 8 km diameter caldera containing a lake The caldera formed about 6800 years ago as a result of the eruption of about 75 km3 of rhyolite magma in the form of tephra, found as far away as Canada, accompanied by pyroclastic flows that left thick deposits of tuff on the flanks of the volcano. Subsequent eruptions have built a cinder cone on the floor of the caldera, which now forms an island called Wizard Island.

8 In stratovolcanoes the collapse and formation of a caldera results from rapid evacuation of the underlying magma chamber by voluminous explosive eruptions that form extensive fall deposits and pyroclastic flows. On shield volcanoes, like in Hawaii, the evacuation of the magma chamber is a slow drawn out processes, wherein magma is withdrawn to erupt on from the rift zones on the flanks. Larger calderas have formed within the past million years in the western United States. These include Yellowstone Caldera in Wyoming, Long Valley Caldera in eastern California, and Valles Caldera in New Mexico.

9 The Yellowstone caldera is an important example, as it illustrates the amount of repose time that might be expected from large rhyolitic systems, and the devastating effect caldera forming eruptions can have on widespread areas. o Yellowstone Caldera which occupies most of Yellowstone National Park, is actually the third caldera to form in the area within the past 2 million years. The three calderas formed at 2.0 million years ago, 1.3 million years ago, and the latest at 600,000 years ago. Thus the repose time is on the average about 650,000 years. o Tephra fall deposits from the latest eruption are found in Louisiana and into the Gulf of Mexico, and covered much of the Western part of the United States. o The eruption 600,000 years ago produced about 1000 km3 of rhyolite (in comparison, the eruption of Mt. St. Helens in May of 1980 produced only 0.75 km3. o Magma still underlies Yellowstone caldera, as evidenced by the large number of hot springs and geysers in the area. Volcanic Eruptions In general, magmas that are generated deep within the Earth begin to rise because they are less dense than the surrounding solid rocks. As they rise they may encounter a depth or pressure where the dissolved gas no longer can be held in solution in the magma, and the gas begins to form a separate phase (i.e. it makes bubbles just like in a bottle of carbonated beverage when the pressure is reduced). When a gas bubble forms, it will also continue to grow in size as pressure is reduced and more of the gas comes out of solution. In other words, the gas bubbles begin to expand. If the liquid part of the magma has a low viscosity, then the gas can expand relatively easily. When the magma reaches the Earth's surface, the gas bubble will simply burst, the gas will easily expand to atmospheric pressure, and a effusive or non-explosive eruption will occur, usually as a lava flow If the liquid part of the magma has a high viscosity, then the gas will not be able to expand very easily, and thus, pressure will build up inside of the gas bubble(s). When this magma reaches the surface, the gas bubbles will have a high pressure inside, which will cause them to burst explosively on reaching atmospheric pressure. This will cause an explosive volcanic eruption

10 and the production of pyroclastic material. Effusive Eruptions Effusive or Non explosive eruptions are favored by low gas content and low viscosity magmas (basaltic to andesitic magmas). If the viscosity is low, non-explosive eruptions usually begin with fire fountains due to release of dissolved gases. Lava flows are produced on the surface, and these run like liquids down slope, along the lowest areas they can find. If the magma emerges along a fracture, it results in a fissure eruption, often called a "curtain of fire" Lava flows produced by eruptions under water are called pillow lavas. If the viscosity is high, but the gas content is low, then the lava will pile up over the vent to produce a lava dome or volcanic dome. Explosive Eruptions Explosive eruptions are favored by high gas content & high viscosity magmas (andesitic to rhyolitic magmas). The explosive bursting of bubbles fragments the magma into clots of liquid that cool as they fall through the air. These solid particles become pyroclasts or volcanic ash. Clouds of gas and tephra that rise above a volcano produce an eruption column that can rise up to 45 km into the atmosphere. Eventually the tephra in the eruption column will be picked up by the wind, carried for some distance, and then fall back to the surface as a tephra fall or ash fall. This type of eruption is called a Plinian eruption. If the pressure in the bubbles is low, the eruption will produce an eruption column only a few hundred meters high, and most of the pyroclastic material will fall to close to the vent to build a cinder cone. This type of eruption is called a Strombolian eruption, and is considered mildly explosive.

11 If the eruption column collapses a pyroclastic flow will occur, wherein gas and tephra rush down the flanks of the volcano at high speed. This is the most dangerous type of volcanic eruption. The deposits that are produced are called ignimbrites if they contain pumice or pyroclastic flow deposits if they contain non-vesicular blocks. A Plinian eruption and pyroclastic flow from Vesuvius volcano killed about 20,000 people in Pompeii in 79 CE. If the gas pressure inside the magma is directed outward instead of upward, a lateral blast can occur. When this occurs on the flanks of a lava dome, a pyroclastic flows called a glowing avalanche or nuée ardentes (in French) can also result. Directed blasts often result from sudden exposure of the magma by a landslide or collapse of a lava dome. This happened at Mt. Pelée Volcano in Martinique in 1902 and killed about 30,000 people. Lahars (Volcanic Mudflows) A volcanic eruption usually leaves lots of loose unconsolidated fragmental debris. When this loose material mixes with water from rainfall, melting of snow or ice, or draining of a crater lake, a mudflow results. Volcanic mudflows are called lahars. These can occur accompanying an eruption or occur long after an eruption. Lahars may be hot or cold and move at high velocity as they fill stream valleys that drain the volcano. At the base of the volcano, they spread out and cover wide areas. In general, they dev estate anything in their path, carrying away homes, buildings, bridges, and destroying roads, and killing livestock and people. In 1985 a lahar produced by a mild eruption of Nevado de Ruiz volcano in Colombia wiped out the village of Armero, about 60 km away from the volcano and killed about 23,000 people.

12 It is important to understand that lahars can occur accompanying an eruption, or can occur simply as the result of heavy rainfall or sudden snow melt, without an eruption. Volcanic Gases Although the predominant gas erupted from volcanoes is H2O vapor, other gases are erupted can have disastrous effects on life. Poisonous gases like Hydrogen Chloride (HCl), Hydrogen Sulfide (H2S), SO2, Hydrogen Fluoride (HF), and Carbon Dioxide (CO2). The Chlorine, Sulfur. and Fluorine gases can kill organisms by direct ingestion, or by absorption onto plants followed by ingestion by organisms. In 1986 an CO2 gas emission from Lake Nyos in Cameroon killed more than 1700 people and 3000 cattle. The gases can also have an effect on the atmosphere and climate. Much of the water on the surface of the earth was produced by volcanoes throughout earth history. Sulfur gases in the atmosphere, along with volcanic ash, reflect incoming solar radiation back into space and have a cooling effect on the atmosphere, thus lowering average global temperatures. The effect lasts only as long as the gases and ash remain in the atmosphere, normally a few years at the most. CO2 gas, produces the opposite effect. It is a greenhouse gas which absorbs solar radiation and causes a warming effect. Eruptions in the past that produced huge quantities of this gas may have been responsible for mass extinction events The Eruption of Mount. St. Helens, 1980 Prior to 1980, Mount St. Helens last erupted in On March 21, 1980 a 4.2 earthquake occurred beneath the volcano signaling the beginning of an eruption. Small eruptions took place through mid April and the summit of the mountain developed a new crater due to the explosions. By the end of April surveys showed that the north face of the mountain had begun

13 to bulge upwards and outwards at rates up to 1 m per day. By May 12, the bulge had displaced parts of the northern part of the volcano a distance of about 150 m. Geologists now recognized that this bulge could soon develop into a landslide. At 8:32 AM on May 18, 1980 a magnitude 5.0 earthquake occurred beneath Mt. St. Helens. This led to a violent eruption that took place over about the next minute. The earthquake triggered a large landslide that began to slide out to the north, initially as three large blocks. As the first block, began to slide downward, the magma chamber beneath the volcano became exposed to atmospheric pressure. The gas inside the magma expanded rapidly, producing a lateral blast that moved outward toward the north. As the second slide block began to move downwards a vertical eruption column began to form above the volcano. The lateral blast rapidly overtook the slide block and roared through an area to the north of the mountain, knocking down all trees in its path and suffocating all living things, Within the next 10 seconds the third slide block moved out toward the north. The landslide thus became a debris avalanche and left a deposit extending about 20 km down the valley (see map below). The southern shores of Spirit Lake were displaced about 1 km northward and the level of the lake was raised about 40 m.

14 Within about the first minute of the eruption the summit of Mount St. Helens had been reduced by about 500 m. The magma however continued to erupt in a Plinian eruption column that reached up to 26 km into the atmosphere. The eruption column collapsed several times to produce pyroclastic flows that moved into Spirit Lake and the upper reaches of the Toutle River Valley. This Plinian phase lasted about 9 hours and spread tephra in a plume to the east, darkening the area at midday to make it appear like night. In all, 62 people lost their lives, either by being buried by the debris avalanche deposit, or suffocating by breathing the hot gases and dust of the blast. Over the next several days melted snow combined with the new ash to produce lahars that roared down the North and South Forks of the Toutle River and drainages to the south of the volcano. In general, the eruption had been much larger than most anticipated, but the fact that a hazards study had been carried out, that public officials were quick to act and evacuate the danger zone, and that the volcano was under constant monitoring, resulted in the minimization of loss of life to only 62 instead of a much larger number that could have been killed had not these efforts been in place. Since the 1980 eruption, several volcanic domes have been emplaced in the crater and some have been blasted out. In the future, it is expected that new domes will continue to form, eventually building the volcano back to a form that will look more like it did prior to the 1980 eruption. Predicting Volcanic Eruptions Before discussing how we can predict volcanic eruptions, its important to get some terminology straight by defining some commonly used terms.

15 Active Volcano - An active volcano to volcanologists is a volcano that has shown eruptive activity within recorded history. Thus an active volcano need not be in eruption to be considered active. Currently there are about 600 volcanoes on Earth considered to be active volcanoes. Each year 50 to 60 of volcanoes actually erupt. Extinct Volcano - An extinct volcano is a volcano that has not shown any historic activity, is usually deeply eroded, and shows no signs of recent activity. How old must a volcano be to be considered extinct depends to a large degree on past activity. Dormant Volcano - A dormant volcano (sleeping volcano) is somewhere between active and extinct. A dormant volcano is one that has not shown eruptive activity within recorded history, but shows geologic evidence of activity within the geologic recent past. Because the lifetime of a volcano may be on the order of a million years, dormant volcanoes can become active volcanoes all of sudden. These are perhaps the most dangerous volcanoes because people living in the vicinity of a dormant volcano may not understand the concept of geologic time, and there is no written record of activity. These people are sometimes difficult to convince when a dormant volcano shows signs of renewed activity. Long - Term Forecasting and Volcanic Hazards Studies Studies of the geologic history of a volcano are generally necessary to make an assessment of the types of hazards posed by the volcano and the frequency at which these types of hazards have occurred in the past. The best way to determine the future behavior of a volcano is by studying its past behavior as revealed in the deposits produced by ancient eruptions. Because volcanoes have such long lifetimes relative to human recorded history, geologic studies are absolutely essential. Once this information is available, geologists can then make forecasts concerning what areas surrounding a volcano would be subject to the various kinds of activity should they occur in a future eruption, and also make forecasts about the long - term likelihood or probability of a volcanic eruption in the area. During such studies, geologists examine sequences of layered deposits and lava flows. Armed with knowledge about the characteristics of deposits left by various types of eruptions, the past behavior of a volcano can be determined. Using radiometric age dating of the deposits the past frequency of events can be determined. This information is then combined with knowledge about the present surface aspects of the volcano to make volcanic hazards maps which can aid other scientists, public officials, and the public at large to plan for evacuations, rescue and recovery in the event that short-term prediction suggests another eruption. Such hazards maps delineate zones of danger expected from the hazards discussed

16 above: lava flows, pyroclastic flows, tephra falls, lahars, floods, etc. Short - Term Prediction based on Volcanic Monitoring Short - term prediction of volcanic eruptions involves monitoring the volcano to determine when magma is approaching the surface and monitoring for precursor events that often signal a forthcoming eruption. Earthquakes - As magma moves toward the surface it usually deforms and fractures rock to generate earthquakes. Thus an increase in earthquake activity immediately below the volcano is usually a sign that an eruption will occur. Ground Deformation - As magma moves into a volcano, the structure may inflate. This will cause deformation of the ground which can be monitored. Instruments like tilt meters measure changes in the angle of the Earth's surface. Other instruments track changes in distance between several points on the ground to monitor deformation. Changes in Heat Flow - Heat is everywhere flowing out of the surface of the Earth. As magma approaches the surface or as the temperature of groundwater increases, the amount of surface heat flow will increase. Although these changes may be small they be measured using infrared remote sensing. Changes in Gas Compositions - The composition of gases emitted from volcanic vents and fumaroles often changes just prior to an eruption. In general, increases in the proportions of hydrogen chloride (HCl) and sulfur dioxide (SO2) are seen to increase relative to the proportion of water vapor. In general, no single event can be used to predict a volcanic eruption, and thus many events are usually monitored so that taken in total, an eruption can often be predicted. Still, each volcano behaves somewhat differently, and until patterns are recognized for an individual volcano, predictions vary in their reliability. Furthermore, sometimes a volcano can erupt with no precursor events at all. Volcanic Hazards The main types of volcanic hazards have been discussed above, so here we only briefly discuss them. You should make sure you understand what each of these are, and what effects each type of hazard can have. We will not likely have time to discuss these again in detail, so the following material is mostly for review. Primary Effects of Volcanism Lava Flows - lava flows are common in Hawaiian and Strombolian type of eruptions, the least explosive. Although they have been known to travel as fast as 64 km/hr, most are slower and give people time to move out of the way. Thus, in general, lava flows are most damaging to property, as they can destroy anything in their path.

17 Pyroclastic Flows - Pyroclastic flows are one of the most dangerous aspects of volcanism. They cause death by suffocation and burning. They can travel so rapidly that few humans can escape. Ash falls - Although tephra falls blanket an area like snow, they are far more destructive because tephra deposits have a density more than twice that of snow and tephra deposits do not melt like snow an cause the collapse of roof. They and can affect areas far from the eruption. Tephra falls destroy vegetation, including crops, and can kill livestock that eat the ash covered vegetation. Tephra falls can cause loss of agricultural activity for years after an eruption. Poisonous Gas Emissions, as discussed above. Secondary and Tertiary Effects of Volcanism Mudflows (Lahars) As discussed above, mudflows can both accompany an eruption and occur many years after an eruption. They are formed when water and loose ash deposits come together and begin to flow. The source of water can be derived by melting of snow or ice during the eruption, emptying of crater lakes during an eruption, or rainfall that takes place any time with no eruption. Debris Avalanches, Landslides, and Debris Flows - Volcanic mountains tend to become oversteepened as a result of the addition of new material over time as well due to inflation of the mountain as magma intrudes. Oversteepened slopes may become unstable, leading to a sudden slope failure that results in landslides, debris flows or debris avalanches. Debris avalanches, landslides, and debris flows do not necessarily occur accompanied by a volcanic eruption. There are documented cases of such occurrences where no new magma has been erupted. Flooding - Drainage systems can become blocked by deposition of pyroclastic flows and lava flows. Such blockage may create a temporary dam that could eventually fill with water and fail resulting in floods downstream from the natural dam. Volcanoes in cold climates can melt snow and glacial ice, rapidly releasing water into the drainage system and possibly causing floods. Jokaulhlaups occur when heating of a glacier results in rapid outburst of water from the melting glacier. Tsunami - Debris avalanche events, landslides, caldera collapse events, and pyroclastic flows entering a body of water may generate tsunami. During the 1883 eruption of Krakatau volcano, in the straits of Sunda between Java and Sumatra, several tsunami were generated by pyroclastic flows entering the sea and by collapse accompanying caldera formation. The tsunami killed about 36,400 people, some as far away from the volcano as 200 km. Volcanic Earthquakes - Earthquakes usually precede and accompany volcanic

18 eruptions, as magma intrudes and moves within the volcano. Although most volcanic earthquakes are small, some are large enough to cause damage in the area immediately surrounding the volcano, and some are large enough to trigger landslides and debris avalanches, such as in the case of Mount St. Helens. Atmospheric Effects- Fined grained ash and sulfur gases expelled into the atmosphere reflect solar radiations and cause cooling of the atmosphere. CO2 released by volcanoes can cause warming of the atmosphere. Volcanoes and Plate Tectonics Global Distribution of Volcanoes In the discussion we had on igneous rocks and how magmas form, we pointed out that since the upper parts of the Earth are solid, special conditions are necessary to form magmas. These special conditions do not exist everywhere beneath the surface, and thus volcanism does not occur everywhere. If we look at the global distribution of volcanoes we see that volcanism occurs four principal settings. 1. Along divergent plate boundaries, such as Oceanic Ridges or spreading centers. 2. In areas of continental extension (that may become divergent plate boundaries in the future). 3. Along converging plate boundaries where subduction is occurring. 4. And, in areas called "hot spots" that are usually located in the interior of plates, away from the plate margins. Since we discussed this in the lecture on igneous rocks, we only briefly review this material here. Diverging Plate Margins Active volcanism is currently taking place along all of oceanic ridges, but most of this volcanism is submarine volcanism. One place where an oceanic ridge reaches above sea level is at Iceland, along the Mid-Atlantic Ridge. Here, most eruptions are basaltic in nature, but, many are explosive strombolian types or explosive phreatic or phreatomagmatic types. As seen in the map to the right, the Mid-Atlantic ridge runs directly through Iceland

19 Volcanism also occurs in continental areas that are undergoing episodes of rifting. A classic example is the East African Rift Valley, where the African plate is being split. The extensional deformation occurs because the underlying mantle is rising from below and stretching the overlying continental crust. Upwelling mantle may melt to produce magmas, which then rise to the surface, often along normal faults produced by the extensional deformation. Basaltic and rhyolitic volcanism is common in these areas. In the same area, the crust has rifted apart along the Red Sea, and the Gulf of Aden to form new oceanic ridges. This may also be the fate of the East African Rift Valley at some time in the future. Other areas where extensional deformation is occurring within the crust is Basin and Range Province of the western U.S. (eastern California, Nevada, Utah, Idaho, western Wyoming and Arizona) and the Rio Grande Rift, New Mexico. These are also areas of recent basaltic and rhyolitic volcanism. Converging Plate Margins All around the Pacific Ocean, is a zone often referred to as the Pacific Ring of Fire, where most of the world's most active and most dangerous volcanoes occur. The Ring of Fire occurs because most of the margins of the Pacific ocean coincide with converging margins along which subduction is occurring

20 The convergent boundary along the coasts of South America, Central America, Mexico, the northwestern U.S. (Northern California, Oregon, & Washington), western Canada, and eastern Alaska, are boundaries along which oceanic lithosphere is being subducted beneath continental lithosphere. This has resulted in the formation of continental volcanic arcs that form the Andes Mountains, the Central American Volcanic Belt, the Mexican Volcanic Belt, the Cascade Range, and the Alaskan volcanic arc. The Aleutian Islands (west of Alaska), the Kurile-Kamchatka Arc, Japan, Philippine Islands, and Marianas Islands, New Zealand, and the Indonesian Islands, along the northern and western margins of the Pacific Ocean are zones where oceanic lithosphere is being subducted beneath oceanic lithosphere. These are all island arcs. o o o o o As discussed previously, the magmas are likely generated by flux melting of the mantle overlying the subduction zone to produce basaltic magmas. Through magmatic differentiation, basaltic magmas change to andesitic and rhyolitic magma. Because these magmas are often gas rich and have all have relatively high viscosity, eruptions in these areas tend to be violent, with common Strombolian, Plinian and Pelean eruptions. Volcanic landforms tend to be cinder cones, stratovolcanoes, volcanic domes, and calderas. Repose periods between eruptions tend to be hundreds to thousands of years, thus giving people living near these volcanoes a false sense of security.

21 Hot Spots Volcanism also occurs in areas that are not associated with plate boundaries, in the interior of plates. These are most commonly associated with what is called a hot spot. Hot spots appear to result from plumes of hot mantle material upwelling toward the surface, independent of the convection cells though to cause plate motion. Hot spots tend to be fixed in position, with the plates moving over the top. As the rising plume of hot mantle moves upward it begins to melt to produce magmas. These magmas then rise to the surface producing a volcano. But, as the plate carrying the volcano moves away from the position over the hot spot, volcanism ceases and new volcano forms in the position now over the hot spot. This tends to produce chains of volcanoes or seamounts (former volcanic islands that have eroded below sea level). The Hawaiian Ridge is one such hot spot trace. Here the Big Island of Hawaii is currently over the hot spot, the other Hawaiian islands still stand above sea level, but volcanism has ceased. Northwest of the Hawaiian Islands, the volcanoes have eroded and are now seamounts. Plateau Basalts or Flood Basalts

22 Plateau or Flood basalts are extremely large volume outpourings of low viscosity basaltic magma from fissure vents. The basalts spread huge areas of relatively low slope and build up plateaus. Many of these outpourings appear to have occurred along a zone that eventually developed into a rift valley and later into a diverging plate boundary. In Oregon and Washington of the northwestern U.S., the Columbia River Basalts represent a series of lava flows all erupted within about 1 million years 12 million years ago. One of the basalt flows, the Roza flow, was erupted over a period of a few weeks traveled about 300 km and has a volume of about 1500 km3. (PREPARED BY GDC HANDWARA)

23 Magmatic Differentiation Differentiation is the process by which magmas evolve to give rise to a variety of magmas and rock types (that have different compositions). Therefore, certain physical processes are required to cause the chemical diversification of a magma (i.e. its differentiation). The chemical trends of magmatic differentiation are often determined by studying crystal liquid relations, but the degree or extent of differentiation is controlled by the efficiency of the differentiation mechanism. In this chapter, we will examine the mechanisms of magmatic differentiation, having to some extent, covered their effects in the last two chapters (crystallization paths: crystal-liquid equilibria, and chemical effects: variation diagrams). We will then briefly discuss the application of trace element, and stable and radiogenic isotope geochemistry in identifying some of the mechanisms of differentiation. Mechanisms of magma diversification (differentiation): 1- Partial melting (to produce different magmas) 2- Crystal fractionation 3- Thermogravitational diffusion 4- Liquid immiscibility 5- Vapor transport 6- Magma mixing 7- Assimilation 1- Partial Melting: (a) Equilibrium partial melting: Partially melting different rock types could be an effective way for producing a variety of magmas of different compositions. A series of liquids produced by successive stages of equilibrium partial melting will upon crystallization produce a sequence of rocks that is the exact opposite of that produced by fractional crystallization. Telling these two mechanisms apart is fairly simple through the use of trace element geochemistry (see later). (b) Fractional melting or incremental batch melting: This is a much more efficient mechanism of differentiation which will depend largely on the frequency at which the liquid is removed from the system. Keep in mind that fractional melting is not a continuous process, and will be arrested for a while as soon as one of the phases is completely used up in the liquid. It will resume once the T is high enough to melt the mixture of the remaining solid phases. The resulting magmas will therefore show sharp differences in composition (as opposed to the gradual compositional changes observed in the case of equilibrium partial melting. (c) Zone melting: Consider a magma chamber undergoing cooling and crystallization from the bottom upward (perhaps because the H2O pressure at the bottom of the chamber is lower than at the top, causing the liquidus T to be suppressed at the top, hence delaying crystallization).

24 As the magma crystallizes at the base, the heat of crystallization released may cause the roof of the magma chamber to partially melt. Accordingly, the melt will appear as if it is migrating upwards. As this happens, elements will be fractionated between the crystals forming at the bottom of the chamber and the melt forming at the top, and between hotter and colder layers of the melt, causing the melt to become more differentiated over time. This fractionation, particularly between hotter and colder liquids, is known as the Soret effect, and is used in refining metals (heavier metals fractionate into colder liquids). This mechanism is unlikely to play a major role in producing large quantities of differentiated magma unless the zone of melting moves very slowly, or the process is repeated many times. It is more efficient and common at depths where the T is sufficiently high. 2- Crystal fractionation: Is the separation of crystals from the melt, either during or after their crystallization. (To me, fractional crystallization implies the continuous removal of crystals from the melt while they are forming, crystal fractionation is a less specific term). Crystal fractionation is one of the most important mechanisms of differentiation that many inexperienced geologists tend to think that both terms (fractionation & differentiation) are synonymous. They are not! Fractionation is a mechanism, whereas differentiation is a phenomenon (or the result of one or more processes). Crystal fractionation is important because of the differences in chemical composition between the crystal and the liquid with which it is in equilibrium, as we have seen in our discussion of crystal - liquid equilibria. However, for crystal fractionation to become an efficient mechanism of differentiation, the whole system has to be somewhat dynamic, with either the crystals moving through the body of the magma, or the magma flowing over a zone of crystallization. The mechanisms of crystal fractionation therefore are: (a) Crystal settling: Minerals crystallizing from a melt may sink to the bottom of magma chambers under the influence of their own weight only if they are denser than the melt. The densities of many mafic minerals, which crystallize at an early stage of cooling (e.g. olivine, chromite, and Opx) are higher than those of the magmas from which they crystallize (see the section on densities of magmas). However, for these minerals to sink, they must overcome the yield strength of the magma. Note that as the magma becomes more acidic, its viscosity (and therefore yield strength) increases, and crystal settling becomes more difficult and unlikely (even though the density of the liquid remaining has decreased). Crystal settling (which to some extent follows Stokes Law), can therefore be an effective way of fractionation in the case of large basic intrusions that cool slowly. Despite its success in explaining some of the phenomena observed in large basic layered intrusions, it cannot be considered the only mechanism of differentiation in these bodies, as is ever so evident in the case of the Skaergaard intrusion! On the other hand, in smaller intrusions where the rate of cooling is high and diffusion is slow, crystal settling has almost no effect on differentiation of the magma. This is because many crystals will form from the melt in these small bodies over a relatively short period of time, and will not be able to sink easily through a mixture of other crystals and liquid.

25 (b) Filter Pressing: If a mixture of crystals and liquid is suddenly subjected to compressional stress, the liquid will be squeezed out of the mixture, and will therefore be separated from the crystals. Although this is a viable mechanism of fractionation, it is not considered to have played a major role in the differentiation of magmas. (c) Flow segregation: Consider the flow of a magma through a fracture. The velocity of flow will be higher in the center of this fracture or dyke to be, than at its edges, where frictional forces are strong (Fig. 1). Accordingly, crystals within this magma will tend to migrate towards the center of the dyke, where they can flow and grow freely, and become coarser grained than at the margins. However, one must be careful in interpreting dykes with coarse-grained cores or centers as indicating flow segregation, since in situ crystallization of magma (w/o flow) in a dyke will result in similar features (cf. chilled margins). Identifying which mechanism is responsible for these observations hinges on identifying mineralogical differences between the centers of dykes and their margins; if such differences are prominent, then flow segregation is a likely cause. If they are more or less the same (save for some contamination), then in situ crystallization is more likely. 3- Thermogravitational Diffusion: Magma chambers may become stratified with different layers having different compositions. Such layering takes place in response to strong thermal and density gradients that develop within the chamber. As T drops towards the top of the chamber (in response to magma cooling from interaction with the overlying country rocks), the top layers will have a tendency to become denser than the magma at the bottom. On the other hand, crystallization of the magma will result in a compositional gradient with magma in the top parts becoming progressively more differentiated and hence tending to be less dense (Fig. 2). As both processes compete, the magma chamber will become layered, with each layer having its own convection cells. Both material and heat will eventually be exchanged between the different layers (hence the term double diffusive convection), further enhancing the compositional differences between the layers, causing the magma to undergo more differentiation. This process is often enhanced by the absorption of water from the roof rocks, and diffusional exchange of the various layers with the wall rocks. Thermogravitational diffusion has been successful at accounting for many observations in layered basic intrusions as well as interlayered acidic tuff sequences (Fig. 3; Please read the figure captions for Figs. 2 & 3 carefully!) 4- Liquid Immiscibility: During crystallization of magmas, and as the composition and T of the liquid changes, this liquid may separate into 2 immiscible liquids. This process is limited to particular magmas with specific compositions, and has been documented in experiments and natural rocks. It applies to: (i) Fe-rich tholeiites which segregate an Fe- and P - rich liquid and a more siliceous one, (ii) some alkaline magmas which segregate a Na + SiO2 rich liquid from a carbonate rich one (ultimately giving rise to carbonatites), and (iii) mafic

26 magmas where a sulfide rich liquid separates from the silicate magma. Figure 4 shows examples of the first two of these 3 cases, and their effects (e.g. formation of brown globules in the mesostasis of a tholeiite, or of a carbonatite magma). 5- Magma Mixing: Mixing two magmas that are compositionally different will produce a magma of intermediate composition (cf. mixing lines on variation diagrams). The effects of this mechanism are most obvious if one magma is basic and the other is acidic, where the basic one will tend to cool and crystallize, while the acidic one will be superheated. For magma mixing to occur, both magmas have to overcome their density contrasts, which will work at separating them into two distinct layers. Several models have been proposed to overcome such density differences, and it is generally considered that blending the 2 magmas becomes much easier if the volume of the basic magma is larger. Sparks et al. (1980) presented a model for the mixing of 2 basic magmas undergoing different degrees of fractional crystallization which will allow them to mix turbulently as their densities (and density contrasts) change (Fig. 5). Magma mixing is more common at the sites of mid-oceanic ridges, where pulses of less differentiated magmas are frequently injected into a fractionated magma in the chamber beneath the ridge. Features that support magma mixing include resorbed xenocrysts (when the phenocrysts of one magma are out of equilibrium with the other mixing magma), and net veined agmatites. Note that the latter is more indicative of magma mixing, as resorbed xenocrysts also form by assimilation. 6- Assimilation: Is the reaction of the magma with the country rocks, whereby these country rocks are incorporated in the magma and eventually melt. For this process to become an efficient mechanism of differentiation, relatively large amounts of the country rocks have to be assimilated by the magma, and/or the compositions of these country rocks have to be drastically different from that of the magma. As in the case of magma mixing, assimilation will produce a magma intermediate in composition between the original magma and the country rock. Assimilation may also result in changes of PH2O of the magma, especially if the assimilated rocks are H2O rich. This will then lower the liquidus T of the magma, and delay crystallization. Assimilation requires thermal energy to heat and possibly melt or partially melt the country rocks, and becomes easier if the assimilated rock is more acidic than the assimilating magma. Otherwise, melting of the assimilated rocks will not occur, and the magma will end up with many xenoliths. For the country rock to melt, a portion of the magma must crystallize and release heat necessary for melting. The amount of melted country rock will always be smaller than that portion of the magma which has crystallized and supplied the necessary heat. Assimilation will therefore be more effective if the magma is rising adiabatically and intersecting its own solidus (thus crystallizing in part and releasing the heat necessary for assimilation).

27 Criteria for recognition of assimilation: 1- The occurrence of xenoliths in the igneous rocks, which are of similar composition to the intruded country rocks. 2- Resorbed xenocrysts (Fig. 6). 3- On variation diagrams, the composition of the igneous rock after assimilation lies on a mixing (straight) line between its composition prior to assimilation and the composition of the assimilated rock. 4- Higher 87Sr/86Sr and 18O values (as will be discussed later). 7- Vapor Phase alteration: During shallow level crystallization of the magma, volatiles may separate from the liquid as they are excluded from the crystallizing phases (when micas and amphiboles become unstable). This phenomenon is termed retrograde boiling, because it may take place during magma cooling. Separation of the volatiles from the magma may be associated with fractionation of some elements between the liquid and vapor phases. Elements that are preferentially incorporated in the vapor phase include Na, K, Si, F, and Cl. This fractionation explains why peralkaline lavas are common, whereas peraluminous ones are rare or absent! The effects of vapor phase transport are most obvious in some tuffs, where the vapor has the ability to metasomatise or alter them. This alteration appears in the form of precipitation of K-feldspar, albite, alkali pyroxenes and amphiboles, and tridymite in the pores of these rocks. A tuff affected by vapor phase alteration is termed a Sillar.

28 Magma Mixing Density & viscosity contrasts Turbulent flow as a means of overcoming density & viscosity contrasts Easier to accomplish between magmas of similar compositions Most common between mantle derived and crustal melts, or at MOR s Example at MOR s Recognition: 1- Net veined agmatites 2- Resorbed phenocrysts (xenocrysts in this case!!!) 3- Reverse zoning in phenocrysts 4- Straight line patterns on variation diagrams (unless ) 5- Isotopic signatures Assimilation Importance of heat of crystallization Need to crystallize 2.5 parts of the magma to assimilate only one part! Theoretically, a magma can assimilate up to 40% of its volume! One of the few processes that can push compositions across thermal divides Recognition 1- xenocrysts 2- xenoliths and schlieren 3- straight line trends on variation diagrams 4- trace element patterns are most strongly affected by assimilation 5- isotopic signatures Vapor phase transport Generation of a vapor phase: (a) release of P during magma ascent (b) fractional crystallization of anhydrous phases (c) assimilation of country rocks Fractionation of elements between fluids and magma: light elements partition preferentially into the vapor phase Effects of release of juvenile fluids: 1- fenitization 2- vapor phase alteration (production of sillars) 3- miarolitic cavities GDC HANDWARA

29 Lecture notes,2nd semester,unit -2 Magma and Igneous Rocks Igneous Rocks are formed by crystallization from a liquid, or magma. They include two types Volcanic or extrusive igneous rocks form when the magma cools and crystallizes on the surface of the Earth Intrusive or plutonic igneous rocks wherein the magma crystallizes at depth in the Earth. Magma is a mixture of liquid rock, crystals, and gas. Characterized by a wide range of chemical compositions, with high temperature, and properties of a liquid. Magmas are less dense than surrounding rocks, and will therefore move upward. If magma makes it to the surface it will erupt and later crystallize to form an extrusive or volcanic rock. If it crystallizes before it reaches the surface it will form an igneous rock at depth called a plutonic or intrusive igneous rock. Types of Magma Chemical composition of magma is controlled by the abundance of elements in the Earth. Si, Al, Fe, Ca, Mg, K, Na, H, and O make up 99.9%. Since oxygen is so abundant, chemical analyses are usually given in terms of oxides. SiO2 is the most abundant oxide. 1. Mafic or Basaltic-- SiO wt%, high in Fe, Mg, Ca, low in K, Na 2. Intermediate or Andesitic-- SiO wt%, intermediate. in Fe,

30 Mg, Ca, Na, K 3. Felsic or Rhyolitic-- SiO %, low in Fe, Mg, Ca, high in K, Na. Gases - At depth in the Earth nearly all magmas contain gas. Gas gives magmas their explosive character, because the gas expands as pressure is reduced. Mostly H2O with some CO2 Minor amounts of Sulfur, Cl, and F Felsic magmas usually have higher gas contents than mafic magmas. Temperature of Magmas Mafic/Basaltic oC Intermediate/Andesitic oC Felsic/Rhyolitic oC. Viscosity of Magmas Viscosity is the resistance to flow (opposite of fluidity). Depends on composition, temperature, & gas content. Higher SiO2 content magmas have higher viscosity than lower SiO2 content magmas Lower Temperature magmas have higher viscosity than higher temperature magmas. Summary Table Magma Type Mafic or Solidified Solidified Chemical Volcanic Plutonic Temperature Viscosity Gas Content Composition Rock Rock Basalt Gabbro SiO Low Low

31 Basaltic Intermediate Andesite Diorite or Andesitic Felsic or Rhyolitic Rhyolite Granite %, high in oc Fe, Mg, Ca, low in K, Na SiO2 %, intermediate o Intermediate Intermediate C in Fe, Mg, Ca, Na, K SiO2 %, low in Fe, Mg, Ca, oc High High high in K, Na Origin of Magma As we have seen the only part of the earth that is liquid is the outer core. But the core is not likely to be the source of magmas because it does not have the right chemical composition. The outer core is mostly Iron, but magmas are silicate liquids. Thus magmas DO NOT COME FROM THE MOLTEN OUTER CORE OF THE EARTH. Thus, since the rest of the earth is solid, in order for magmas to form, some part of the earth must get hot enough to melt the rocks present. We know that temperature increases with depth in the earth along the geothermal gradient. The earth is hot inside due to heat left over from the original accretion process, due to heat released by sinking of materials to form the core, and due to heat released by the decay of radioactive elements in the earth. Under normal conditions, the geothermal gradient is not high enough to melt rocks, and thus with the exception of the outer core, most of the Earth is solid. Thus, magmas form only under special circumstances. To understand this we must

32 first look at how rocks and mineral melt. As pressure increases in the Earth, the melting temperature changes as well. For pure minerals, there are two general cases. For a pure dry (no H2O or CO2 present) mineral, the melting temperate increases with increasing pressure. For a mineral with H2O or CO2 present, the melting temperature first decreases with increasing pressure Since rocks mixtures of minerals, they behave somewhat differently. Unlike minerals, rocks do not melt at a single temperature, but instead melt over a range of temperatures. Thus, it is possible to have partial melts from which the liquid portion might be extracted to form magma. The two general cases are:

33 Melting of dry rocks is similar to melting of dry minerals, melting temperatures increase with increasing pressure, except there is a range of temperature over which there exists a partial melt. The degree of partial melting can range from 0 to 100% Melting of rocks containing water or carbon dioxide is similar to melting of wet minerals, melting temperatures initially decrease with increasing pressure, except there is a range of temperature over which there exists a partial melt. Three ways to Generate Magmas From the above we can conclude that in order to generate a magma in the solid part of the earth either the geothermal gradient must be raised in some way or the melting temperature of the rocks must be lowered in some way. The geothermal gradient can be raised by upwelling of hot material from below either by uprise solid material (decompression melting) or by intrusion of magma (heat transfer). Lowering the melting temperature can be achieved by adding water or Carbon Dioxide.

34 Decompression Melting - Under normal conditions the temperature in the Earth, shown by the geothermal gradient, is lower than the beginning of melting of the mantle. Thus in order for the mantle to melt there has to be a mechanism to raise the geothermal gradient. Once such mechanism is convection, wherein hot mantle material rises to lower pressure or depth, carrying its heat with it. If the raised geothermal gradient becomes higher than the initial melting temperature at any pressure, then a partial melt will form. Liquid from this partial melt can be separated from the remaining crystals because, in general, liquids have a lower density than solids. Basaltic magmas appear to originate in this way. Upwelling mantle appears to occur beneath oceanic ridges, at hot spots, and beneath continental rift valleys. Thus, generation of magma in these three environments is likely caused by decompression melting. Transfer of Heat- When magmas that were generated by some other mechanism intrude into cold crust, they bring with them heat. Upon

35 solidification they lose this heat and transfer it to the surrounding crust. Repeated intrusions can transfer enough heat to increase the local geothermal gradient and cause melting of the surrounding rock to generate new magmas. Transfer of heat by this mechanism may be responsible for generating some magmas in continental rift valleys, hot spots, and subduction related environments. Flux Melting - As we saw above, if water or carbon dioxide are added to rock, the melting temperature is lowered. If the addition of water or carbon dioxide takes place deep in the earth where the temperature is already high, the lowering of melting temperature could cause the rock to partially melt to generate magma. One place where water could be introduced is at subduction zones. Here, water present in the pore spaces of the subducting sea floor or water present in minerals like hornblende, biotite, or clay minerals would be released by the rising temperature and then move in to the overlying mantle. Introduction of this water in the mantle would then lower the melting temperature of the mantle to generate partial melts, which could then separate from the solid mantle and rise toward the surface.

36 Chemical Variability of Magmas The chemical composition of magma can vary depending on the rock that initially melts (the source rock), and process that occur during partial melting and transport. Initial Composition of Magma The initial composition of the magma is dictated by the composition of the source rock and the degree of partial melting. In general, melting of a mantle source (garnet peridotite) results in mafic/basaltic magmas. Melting of crustal sources yields more siliceous magmas. In general more siliceous magmas form by low degrees of partial melting. As the degree of partial melting increases, less siliceous compositions can be generated. So, melting a mafic source thus yields a felsic or intermediate magma. Melting of ultramafic (peridotite source) yields a basaltic magma. Magmatic Differentiation But, processes that operate during transportation toward the surface or during storage in the crust can alter the chemical composition of the magma. These processes are referred to as magmatic differentiation and include assimilation, mixing, and fractional crystallization. Assimilation - As magma passes through cooler rock on its way to the surface it may partially melt the surrounding rock and incorporate this melt into the magma. Because small amounts of partial melting result in siliceous liquid compositions, addition of this melt to the magma will make it more siliceous. Mixing - If two magmas with different compositions happen to come in contact with one another, they could mix together. The mixed magma will have a composition somewhere between that of the original two magma compositions. Evidence for mixing is often

37 preserved in the resulting rocks. Fractional Crystallization - When magma crystallizes it does so over a range of temperature. Each mineral begins to crystallize at a different temperature, and if these minerals are somehow removed from the liquid, the liquid composition will change. The processes is called magmatic differentiation by Fractional Crystallization. Because mafic minerals like olivine and pyroxene crystallize first, the process results in removing Mg, Fe, and Ca, and enriching the liquid in silica. Thus crystal fractionation can change a mafic magma into a felsic magma. Crystals can be removed by a variety of processes. If the crystals are more dense than the liquid, they may sink. If they are less dense than the liquid they will float. If liquid is squeezed out by pressure, then crystals will be left behind. Removal of crystals can thus change the composition of the liquid portion of the magma. Let me illustrate this using a very simple case. Imagine a liquid containing 5 molecules of MgO and 5 molecules of SiO2. Initially the composition of this magma is expressed as 50% SiO2 and 50% MgO. i.e. Now let's imagine I remove 1 MgO molecule by putting it into a crystal and removing the crystal from the magma. Now what are the percentages of each molecule in the liquid? If we continue the process one more time by removing one more MgO

38 molecule Thus, composition of liquid can be changed. Bowen's Reaction Series Bowen found by experiment that the order in which minerals crystallize from a basaltic magma depends on temperature. As a basaltic magma is cooled Olivine and Ca-rich plagioclase crystallize first. Upon further cooling, Olivine reacts with the liquid to produce pyroxene and Ca-rich plagioclase react with the liquid to produce less Ca-rich plagioclase. But, if the olivine and Ca-rich plagioclase are removed from the liquid by crystal fractionation, then the remaining liquid will be more SiO2 rich. If the process continues, an original basaltic magma can change to first an andesite magma then a rhyolite magma with falling temperature Distribution of Igneous Activity Igneous activity is currently taking place as it has in the past in various tectonic settings. These include diverging and converging plate boundaries, hot spots, and rift valleys. Divergent Plate Boundaries At oceanic ridges, igneous activity involves eruption of basaltic lava flows that form pillow lavas at the oceanic ridges and intrusion of dikes and plutons beneath the ridges. The lava flows and dikes are basaltic and the plutons mainly gabbros. These processes form the bulk of the oceanic crust as a result of sea floor spreading. Magmas are generated by decompression melting as hot solid asthenosphere rises and partially

39 melts. Convergent Plate Boundaries Subduction at convergent plate boundaries introduces water into the mantle above the subduction and causes flux melting of the mantle to produce basaltic magmas. These rise toward the surface differentiating by assimilation and crystal fractionation to produce andesitic and rhyolitic magmas. The magmas that reach the surface build island arcs and continental margin volcanic arcs built of basalt, andesite, and rhyolite lava flows and pyroclastic material. The magmas that intrude beneath these arcs can cause crustal melting and form plutons and batholiths of diorite and granite Hot Spots As discussed previously, hot spots are places are places where hot mantle ascends toward the surface as plumes of hot rock. Decompression melting in these rising plumes results in the production of magmas which erupt to form a volcano on the surface or sea floor, eventually building a volcanic island. As the overriding plate moves over the hot spot, the volcano moves off of the hot spot and a new volcano forms over the hot spot. This produces a hot spot track consisting of lines of extinct volcanoes leading to the active volcano at the hot spot. A hot spot located beneath a continent can result in heat transfer melting of the continental crust to produce large rhyolitic volcanic centers and plutonic granitic plutons below. A good example of a continental hot spot is at Yellowstone in the western U.S. Occasionally a hot spot is coincident with an oceanic ridge. In such a case, the hot spot produces larger volumes of magma than normally occur at ridge and thus build a volcanic island on the ridge. Such is the case for Iceland which sits atop the Mid-Atlantic Ridge. Rift Valleys Rising mantle beneath a continent can result in extensional fractures in

40 the continental crust to form a rift valley. As the mantle rises it undergoes partial melting by decompression, resulting in the production of basaltic magmas which may erupt as flood basalts on the surface. Melts that get trapped in the crust can release heat resulting in melting of the crust to form rhyolitic magmas that can also erupt at the surface in the rift valley. An excellent example of a continental rift valley is the East African Rift. Large Igneous Provinces In the past, large volumes of mostly basaltic magma have erupted on the sea floor to form large volcanic plateaus, such as the Ontong Java Plateau in the eastern Pacific. Such large volume eruptions can have affects on the oceans because they change the shape of ocean floor and cause a rise in sea level, that sometimes floods the continents. The plateaus form obstructions which can drastically change ocean currents. These changes in the ocean along with massive amounts of gas released by the magmas can alter climate and have drastic effects on life on the planet. (COMPILED BY GDC HANDWARA)

41 1 Magmas and Lavas Definitions: Magma: a mixture of a melt (predominantly silicate) ± crystals ± volatiles which occurs at depths and has the ability to migrate to shallower levels where it either crystallizes at depth giving rise to igneous intrusions, or erupts at the surface to form volcanic rocks. Lava: Is erupted molten material that can flow on the surface of the earth. A lava may therefore be considered a magma that has lost its gases (to the atmosphere upon eruption). Chemical composition of magmas (types of magmas): It is clear from our discussion of the layered structure of the earth that although magmas are generated by partial melting in the upper mantle or lower crust, such a process occurs over a range of depths. Accordingly, not all magmas have the same composition. This is evidenced by the variety of igneous rocks that occur at the surface of the earth or at depth. Volcanic eruptions also show that lavas have different viscosities, which in turn are due to their different chemical compositions and/or temperatures. By carefully studying the chemistry of the different types of igneous rocks, and their associations with each other, petrologists were able to classify magmas into four main chemical groups: 1- Acidic: rich in SiO2, Na2O and K2O. Rocks produced from such magmas have between 66 and 77.5% by weight SiO2. "Granite" is an example of an acidic rock, and many acidic magmas are broadly known as "granitic". 2- Intermediate: rich in SiO2, Na2O, K2O as well as CaO and Al2O3. Rocks produced from such magmas have SiO2 values in the range 52 to 66% by weight. Andesite is a good example of a rock formed by the crystallization of an intermediate lava. 3- Basic: rich in CaO, MgO and FeO. Rocks of this type have SiO2 values of 45-52% by weight. Basalt is an example of a basic rock, and many basic magmas are broadly known as "basaltic". 4- Ultrabasic: Are magmas poor in SiO2 (< 45%) but with large amounts of FeO and MgO. Ultrabasic rocks may have SiO2 values as low as 38% by weight. Peridotite is a good example of an ultrabasic rock. Note that ultrabasic lavas are almost nonexistent, being restricted to Precambrian terranes! (Can you guess why?) Table 1 lists the chemical compositions of some igneous rocks belonging to these four types. Technically speaking, a rock can only be considered basic, acidic or intermediate if its chemical composition (particularly its SiO2 content) is known. In the absence of such information, it is better to use the more general and descriptive terms: "ultramafic, mafic and felsic" which are based on the mineralogy of the rock. The term ultramafic is used to describe rocks that have very little or no feldspar, and which are very rich in darkcoloured ferromagnesian minerals as olivine and pyroxene. Mafic is used for rocks rich in ferromagnesian minerals, but which also contain some feldspars. Felsic rocks are those rich in feldspars and quartz. Because ultrabasic and basic rocks are dark in colour and are rich in Fe and Mg, whereas acidic rocks tend to be light in colour and rich in quartz and feldspars, the terms ultrabasic and ultramafic on one hand and acidic and felsic on the other have unfortunately been used interchangeably.

42 2 Table 1: Average chemical compositions of selected igneous rock types Oxide SiO2 TiO2 Al2O3 Fe2O3 FeO MnO MgO CaO Na2O K2O H2O CO2 P2O5 Acidic (Granite) Intermediate (Andesite) Basic (Basalt) Ultrabasic (Peridotite) In the course of studying igneous rocks, it was found that acidic rocks melt at lower temperature compared to basic and ultrabasic rocks. Because melting is the reverse of crystallization, understanding how a rock melts will help us understand the process of magma formation, as well as the process of formation of igneous rocks from a magma. Formation of magmas: The temperature in the earth generally increases regularly with increasing depth, and the variation of temperature with depth in a specific area and at a specific time in the earth's history is known as the "geotherm", the slope of which is known as the "geothermal gradient". Different areas have different geothermal gradients (Fig. 1) which may also change over time (as the tectonic setting of that area changes). As you are aware, the highest geotherms occur beneath mid-oceanic ridges (Fig. 1). If a rock is buried deep below the surface, it will first be metamorphosed, then if the melting point of some of its constituent minerals is reached, the rock will begin to melt. Keep in mind that the melting temperatures of minerals will change as a function of pressure along a curve known as the melting curve. Because different minerals have different melting curves, and because a rock is an aggregate of different minerals, melting of a rock will take place over a range of temperatures bounded by two curves: the solidus and the liquidus (Fig. 2). This is the process known as "partial melting", since only part of the rock melts at any given temperature. The solidus is the curve of all temperatures below which the rock is completely in the solid state. The liquidus is the curve joining all temperatures above which the rock is totally molten. Between the two curves, solids and a melt coexist in equilibrium.

43 3 In order to achieve any degree of melting of a particular rock within the crust or mantle, temperatures must exceed those defined by the solidus of that rock. Figure 3 shows the relationship between two different geotherms (one beneath the continents, and the other beneath the oceans), and the melting curves of an acidic rock (a granite, in the presence of H2O; Fig. 3b) and that of an ultrabasic rock (a dry peridotite; Fig. 3a). In either case, melting can only occur at depths greater than those defined by the intersection of the geotherms with the solidus for that rock (shaded areas). It is clear from this figure that because of the different melting temperatures of a wet granite and a peridotite, an acidic melt can be generated at depths as low as 35 km, whereas a basic magma may require depths of >300 km to form! In order to understand how a magma is actually generated in the mantle, we need to know what kind of rocks occur in the mantle. We have already decided that most magmas are produced in layers (a) - (d) of the mantle (collectively known as the upper mantle), and that only small amounts of magmas form in the lower crust. The upper mantle layers consist of different types of peridotite. In the uppermost parts of the lithospheric mantle (especially beneath the oceanic crust), the pressures are low enough for plagioclase to be stable, and the peridotites at these depths may contain minor amounts of plagioclase, in addition to the common Opx, Cpx and olivine. Such peridotites are therefore called plagioclase lherzolites. At greater depths, plagioclase becomes unstable, and the mantle consists predominantly of spinel lherzolites. At greater depths, garnet becomes a common constituent of peridotites which are then known as garnet lherzolites. These relations are shown on Fig. 2. Because different layers of the mantle have different compositions, partial melting at different depths in the mantle will produce different types of magma. However, Fig. 2 also shows that an average geotherm does not intersect the solidus of mantle rocks, suggesting that under normal geothermal conditions, it is not possible to produce a magma from the mantle by partial melting! Partial melting in the mantle will therefore take place if: (a) the geotherm is perturbed, shifting to higher temperatures. (b) H2O or CO2 are added to the mantle, thus shifting their solidi to lower T (Fig. 4) (c) Adiabatic decompression: This model suggests that packets of mantle material can rise very quickly from the deep parts of the mantle into shallower levels while still maintaining their temperature. This causes these packets to cross the solidus of mantle lherzolites, and thus undergo partial melting. A good example of such "packets" is rendered by mantle plumes. Convection in the mantle is believed to be the driving force behind the adiabatic rise of mantle material. Figure 5 shows the process of adiabatic decompression. Once the mantle undergoes partial melting by about 20% or more, the melt fraction (or primary magma) is squeezed out of it, and begins its journey towards the surface of the earth (Fig. 6). Keep in mind that the mantle layers are not necessarily homogeneous, and that they may contain (in addition to basaltic magma) eclogites (Gt + Cpx), harzburgites (Opx + Ol) and dunites (monomineralic rocks of olivine). The latter two rock types are believed to represent what is left behind after partial melting. Accordingly, if the mantle is found to contain considerable amounts of harzburgite + dunite, it is often termed

44 4 "depleted mantle" meaning a mantle that has undergone significant partial melting. Figure 7 shows the relationship between the composition of the magma generated by partial melting, and the depth at which such magmas are generated. Structure of silicate liquids By definition, a magma consists of a melt ± suspended crystals ± dissolved or exsolved gases. This shows that the melt (usually a silicate liquid) is the main constituent of a magma. It is therefore important to understand its structure, which will in turn help us understand the physical properties of magmas. Silicate melts consist of polymers of interconnected by distorted Si-O tetrahedra, with other major ions such as Na, K, Al, Ca, Fe and Mg occurring in looser coordination with O (Fig. 8). In general, elements can be grouped into two main groups as far as their effect on the degree of polymerization of the silicate liquid: (a) Network formers: Si, Al, Na, K, Rb, Cs, Ti and P. Al3+ and a monovalent ion can substitute for Si4+, hence the network forming capability of these elements. Ti and P have large charges (4 and 5, respectively) which allow them to act as network formers. (b) Network modifiers: Fe, Mg, Ca and Mn, all of which substitute into the melt structure by breaking the Si - O bonds. Note that in alkalic magmas in which Na + K > Al, Na and K become network modifiers instead of network formers. Similarly, in highly aluminous melts, Al will tend to occupy the octahedral (instead of the tetrahedral) sites, and will also become a network modifier. H2O, Cl, F, and S are all network modifiers. H2O being the most common volatile in magmas plays a major role in breaking Si-O bonds and lowering the viscosity of the melt (Fig. 9). CO2 on the other hand appears to be a network former, binding broken Si-O bonds back together, although its role is not exactly known. Volatiles in magmas Volatiles in magmas are either dissolved (at high pressures) or exsolved (closer to the surface, where the pressure buildup may result in the violent eruptions of some volcanoes). Experimental studies have shown that magmas can dissolve more gases than have actually been observed. However, determination of the exact amount or composition of volatiles contained in a magma is very difficult, as magmas tend to lose volatiles to the atmosphere on eruption, or to the intruded country rocks during their slow cooling. Nevertheless, methods of estimating the nature and amount of volatile constituents of a magma include: i) analysis of the tops of pillow lavas, and counting their vesicles. ii) analysis of melt inclusions in crystals iii) analysis of quenched magmas iv) phase equilibria experiments.

45 5 In addition, volcanic eruptions give us some idea of the amount and composition of volcanic gases. Table 1 lists the composition of volcanic gases. All of these studies show that magmas have ~ 1% by weight volatiles, and that H2O is the most common volatile constituent of acidic magmas, whereas CO2 is more common in basaltic ones. Although this may seem like a very small amount, the low molecular weight of volatiles ensures that they will occupy a significantly larger % of the magma by volume. The decrease of the H2O/CO2 ratio as the magmas become more basaltic or more alkalic is due to the fact that CO2 and S gases are more soluble in basaltic magmas. On the other hand, F and Cl are more soluble in acidic magmas. Because H2O is the most common volatile constituent of the magma, it is important to understand the form in which it enters (and dissolves in) a magma. Figure 9 shows that H2O breaks Si-O bonds and is therefore one of the most important network modifiers. Stopler (1982) has shown that at low concentrations of H2O in the magma, this H2O is in the form of OH-; at higher concentrations, H2O enters the magma in its molecular form (Fig. 10). Factors affecting the solubility of gases in magmas: i) pressure: the higher the P, the larger the wt% of H2O dissolved (Fig. 11a). ii) temperature: the higher the T, the lower the wt% of H2O dissolved (Fig. 11b). iii) chemical composition of the magma: acidic magmas contain a larger amount of volatiles (Fig. 11a). This is because such magmas have already undergone a significant amount of crystallization of anhydrous phases, thus concentrating the volatiles. iv) composition of the volatiles: An increased amount of CO2 in the volatiles decreases the solubility of H2O in the magma (Fig. 11c). Physical properties of magmas 1) Temperature Temperatures of magmas can be inferred from: i) direct measurement of the temperature of lavas using optical pyrometers ii) direct measurement of lava temperatures using thermocouples iii) phase equilibria experiments iv) application of well calibrated geothermometers. Direct measurement methods yield temperatures in the range: C. The lower temperatures represent those determined for acidic rocks or partially crystalline lavas, whereas the higher ones are for basalts. Because most solidi have positive slopes (i.e. the meting temperature increases by ~ 3 C/km, the original temperatures of a magma generated at a depth of ~ 50 km will be 150 C higher than that determined for the corresponding lava at the surface.

46 6 2) Thermal conductivity, specific heat and heat of fusion: The thermal conductivity of both rocks and magmas is very low. It generally increases with increasing temperature. Acidic magmas have a higher thermal conductivity compared to basic ones. Fig. 12 shows the relationship between thermal conductivity and T for several types of magmas. The specific heat (or heat capacity Cp), defined as the heat necessary to raise the temperature of 1 gm of a substance by 1 C, is very low for magmas (~0.3 cal/gm). On the other hand, the heat of fusion of igneous rocks (or the heat of crystallization of a magma, defined as the heat released during the crystallization of 1 gm of a melt) are relatively large (~ cal/gm). This means that, as the magma is crystallizing, it releases a large amount of heat. Because of the low Cp of the same magma, the heat released during crystallization will to a certain extent "raise" the temperature of the magma while the same magma is "cooling" and crystallizing. The overall effect is to maintain the temperature of the magma over a long period of time (i.e. cooling during crystallization will be slow). However, as soon as the magma crystallizes, the igneous rock will cool much more rapidly. 3) Viscosity Viscosity is a measure of the resistance to flow of any fluid substance. Fluids are broadly grouped into two groups according to their viscosity: (i) Newtonian fluids: those which respond directly by flow to any kind of applied shear stress, and (ii) Bingham fluids: those which have a finite yield strength, and will not flow unless this strength is "overcome" by the applied shear stress. Magmas are Bingham fluids. Factors affecting the viscosity of a magma: i) temperature: the higher the T, the lower the viscosity (Fig. 13a & c) ii) pressure: has a small effect on viscosity; the higher the pressure, the lower the viscosity (Fig. 13d). iii) composition: Acidic magmas are more viscous than basic ones. This is simply a function of the percentage of network formers and network modifiers in each magma; the higher SiO2 content of acidic magmas ensures that they have the larger % of network formers. This relationship is shown in Fig. 13a. iv) amount and nature of volatiles: Again this is a reflection of the ability of volatile components to act as network modifiers or (in case of CO2) as network formers. Addition of a small amount of H2O to a magma drastically lowers its viscosity (Fig. 13c), especially if this magma is acidic or intermediate. On the other hand, addition of H2O to an olivine melt will only have a minor effect on its viscosity. This is because H2O breaks Si-Al-O bonds, and Olivine has no Al! Viscosity of a magma largely controls the structures formed upon its crystallization. For example, fissure eruptions or shield volcanoes form from basaltic lavas which have low viscosities, whereas cumulo-domes and plugs are formed from high viscosity acidic lavas.

47 7 4) Density The density of a magma is one of the most important factors controlling its physical and chemical behavior. Density plays a role in controlling the movement of magmas to shallower levels. Moreover, density contrast between a silicate melt and any minerals that may have crystallized from it plays a role in the differentiation of this magma. Magma densities range between 2.2 and 3.1 g/cm3 Mafic minerals such as pyroxenes, olivines and amphiboles are usually denser than the magma, and have a tendency to sink to the bottom of the magma chamber, resulting in a change in the chemical composition of the remaining magma. On the other hand, only calcic plagioclases may be denser than some magmas, but sodic plagioclases and nepheline usually float on the top of all magma. The velocity at which a mineral sinks to the floor of a magma chamber can be estimated by Stokes' law: v = 2 g r2( 1-2)/9 where v is the settling velocity, r is the radius of the settling crystal (assuming it has a spherical shape), ( 1-2) is the density contrast between the magma and the crystal, and is the viscosity of the magma. However, keep in mind that for a crystal to sink, the gravitational force acting on it must be able to exceed the yield strength of the magma. Factors affecting the density of a magma: i) temperature: the higher the T, the lower the density (Fig. 14a) ii) pressure: the higher the pressure, the higher the density. Accordingly, the density of a magma will increase with increasing depth, until this magma is actually denser than the solid rocks (this is because fluids are more compressible than solids). This happens at depths of ~ 400 km. Accordingly, magmas generated from depths > 400 km do not rise to the surface! iii) composition: The density of a magma depends largely on its Fe content. Accordingly, basic magmas are denser than acidic ones (Fig. 14a & b). This plays a major role in differentiation (Fig. 14b). Movement of magmas: Now that we understand how a magma is generated, and have some knowledge of its physical and chemical properties, we need to understand how it moves. Movement of magmas is triggered essentially by density contrast between it and the surrounding or overlying rocks, and is controlled to a large extent by its viscosity. As soon as a melt is formed in lower crust or upper mantle (at depths < 400 km), it will have a tendency to rise to shallower depths by virtue of its lower density. This movement is known as "diapiric rise" (Fig. 14). Movement of magma will also take place through connected pores and channelways, essentially following the laws of fluid flow through porous media. At shallow depths, magmas migrate along fissures and fractures. As the magma moves, it essentially creates new channelways for itself by engulfing country rocks, and possibly melting or reacting with them (a process known as assimilation). The whole process describing the movement of magmas at shallow depths by engulfing country rocks is known as "stoping".

48 8 While the magma is migrating to shallower levels, it loses heat, and begins to crystallize. The processes of differentiation (which include assimilation) also cause the magma to change its composition. Such compositional changes will change the density and viscosity of the magma, and hence affect its ability to continue its journey to the surface. For example, crystallization of 40-50% of the magma (with the crystals remaining suspended in it) will increase its viscosity so much that its rate of flow will be too slow, and it will probably never reach the surface. It is therefore clear that lavas on the surface of the earth are not representative of the original magmas from which they form, and that the lithosphere acts as a barrier that filters out magmas that are denser than its constituent rocks or too viscous to flow through its narrow channelways. Magmas that do not reach the surface, will end up crystallizing at depths as intrusive rocks. Before crystallizing, such magmas pile up in one place referred to as a magma chamber, where they begin to lose heat (± some volatiles??) to the surrounding country rocks. The loss of heat may not be homogeneous throughout the chamber, which in turn causes different parts of the chamber to have different temperatures (i.e. results in the development of thermal gradients within the chamber). The onset of crystallization will have a similar effect, and the possible removal of crystals from parts of the magma (or assimilation of country rocks from the roof of the magma chamber) may also cause the chamber to develop compositional gradients. Either one of these processes may cause the magma to convect by developing simple convection cells (Fig. 16a). Both thermal and compositional gradients operating simultaneously may cause the development of complex convection cells, a process called double diffusive convection (Fig. 16b - d). We will examine the effects of these processes in more detail when we discuss the process of differentiation. GDC HANDWARA

49 Lecture notes,2nd semester unit-2 MAGMATIC DIFFERENTIATION (Chemical Variation in Rock Suites) Soon after geologists began doing chemical analyses of igneous rocks they realized that rocks emplaced in any given restricted area during a short amount of geologic time were likely related to the same magmatic event. Evidence for some kind of relationship between the rocks, and therefore between the magmas that cooled to form the rocks came from plotting variation diagrams. A variation diagram is a plot showing how each oxide component in a rock varies with some other oxide component. Because SiO2 usually shows the most variation in any given suite of rocks, most variation diagrams plot the other oxides against SiO2 as shown in the diagram here, although any other oxide could be chosen for plotting on the x-axis. Plots that show relatively smooth trends of variation of the components suggested that the rocks might be related to one another through some process. Of course, in order for the magmas to be related to one another, they must also have been intruded or erupted within a reasonable range of time. Plotting rocks of Precambrian age along with those of Tertiary age may show smooth variation, but it is unlikely that the magmas were related to one another. If magmas are related to each other by some processes, that process would have to be one that causes magma composition to change. Any process that causes magma composition to change is called magmatic differentiation. Over the years, various process have been suggested to explain the variation of magma compositions observed within small regions. Among the processes are: Distinct melting events from distinct sources. Various degrees of partial melting from the same source. Crystal fractionation. Mixing of 2 or more magmas. Assimilation/contamination of magmas by crustal rocks. Liquid Immiscibility.

50 Initially, researchers attempted to show that one or the other of these process acted exclusively to cause magmatic differentiation. With historical perspective, we now realize that if any of them are possible, then any or all of these processes could act at the same time to produce chemical change, and thus combinations of these processes are possible. Still, we will look at each one in turn in the following discussion. Distinct Melting Events One possibility that always exists is that the magmas are not related except by some heating event that caused melting. In such a case each magma might represent melting of a different source rock at different times during the heating event. If this were the case, we might not expect the chemical analyses of the rocks produced to show smooth trends on variation diagrams. But, because variation diagrams are based on a closed set of numbers (chemical analyses add up to 100%), if the weight% of one component increases, then the weight percent of some other component must decrease. Thus, even in the event that the magmas are not related, SiO2 could increase and MgO could decrease to produce a trend. The possibility of distinct melting events is not easy to prove or disprove. Various Degrees of Partial Melting We have seen in our study of simple phase diagrams that when a multicomponent rock system melts, unless it has the composition of the eutectic, it melts over a range of temperatures at any given pressure, and during this melting, the liquid composition changes. Thus, a wide variety of liquid compositions could be made by various degrees of partial melting of the same source rock. To see this, lets look at a simple example of a three component system containing natural minerals, the system Fo - Di - SiO2, shown in simplified form here. A proxy for mantle peridotite, being a mixture of Ol, Cpx, and Opx would plot as shown in the diagram. This rock would begin to melt at the peritectic point, where Di, En, Ol, and Liquid are in equilibrium. The composition of the liquid would remain at the peritectic point (labeled 0% melting) until all of the diopside melted. This would occur after about 23% melting. The liquid would then take a path shown by the dark curve, first moving along the En - Ol boundary curve, until the enstatite was completely absorbed, then moving in a direct path toward the peridotite composition. At 100% melting the liquid would have the composition of the initial peridotite. So long as some of the liquid is left behind, liquids can be extracted at any time during the melting event and have compositions anywhere along the dark like between 0% melting and 100% melting. (Note that the compositions between 0% melting and where the dark line intersects the En-Di join are SiO2 oversaturated liquids, and those from this point up to 100% melting are SiO2 undersaturated liquids).

51 Fractional Melting Note that it was stated above that some of the liquid must be left behind. If all of the liquid is removed, then we have the case of fractional melting, which is somewhat different. In fractional melting all of the liquid is removed at each stage of the process. Let's imagine that we melt the same peridotite again, removing liquids as they form. The first melt to form again will have a composition of the peritectic, labeled "Melt 1" in the diagram. Liquids of composition - Melt 1 can be produced and extracted until all of the Diopside is used up. At this point, there is no liquid, since it has been removed or fractionated, so the remaining solid consists only of Enstatite and Forsterite with composition "Solid 2". This is a two component system. Thus further melting cannot take place until the temperature is raised to the peritectic temperature in the two component system FoSiO2. Melting at this temperature produces a liquid of composition "Melt 2". Further melting and removal of this liquid, eventually results in all of the Enstatite being used up. At this point, all that is left in the rock is Forsterite. Forsterite melts at a much higher temperature, so further melting cannot take place until the temperature reaches the melting temperature of pure Forsterite. This liquid will have the same composition as pure Forsterite ("Melt 3"). We saw in our discussion of how magmas are generated that it is difficult enough to get the temperature in the Earth above the peridotite solidus, let alone to much higher temperatures. Thus, fractional melting is not very likely to occur in the Earth. Trace Elements as Clues to Suites Produced by Various Degrees of Melting Trace elements are elements that occur in low concentrations in rocks, usually less than 0.1 % (usually reported in units of parts per million, ppm). When considering the rocks in the mantle, trace elements can be divided into incompatible elements, those that do not easily fit into the crystal structure of minerals in the mantle, and compatible elements, those that do fit easily into the crystal structure of minerals in the mantle. Incompatible elements - these are elements like K, Rb, Cs, Ta, Nb, U, Th, Y, Hf, Zr, and the Rare Earth Elements (REE)- La, Ce, Pr, Nd, Pm, Sm, Eu, Gd, Tb, Dy, Ho, Er, Tm, Yb, & Lu. Most have a large ionic radius. Mantle minerals like olivine, pyroxene, spinel, & garnet do not have crystallographic sites for large ions. Compatible elements - these are elements like Ni, Cr, Co, V, and Sc, which have smaller ionic radii and fit more easily into crystallographic sites that normally accommodate Mg, and Fe.

52 When a mantle rock begins to melt, the incompatible elements will be ejected preferentially from the solid and enter the liquid. This is because if these elements are present in minerals in the rock, they will not be in energetically favorable sites in the crystals. Thus, a low degree melt of a mantle rock will have high concentrations of incompatible elements. As melting proceeds the concentration of these incompatible elements will decrease because (1) there will be less of them to enter the melt, and (2) their concentrations will become more and more diluted as other elements enter the melt. Thus, incompatible element concentrations will decrease with increasing % melting. Rare Earth elements are particularly useful in this regard. These elements, with the exception of Eu, have a +3 charge, but their ionic radii decrease with increasing atomic number. i.e. La is largest, Lu is smallest. Thus the degree of incompatibility decreases from La to Lu. This is even more true if garnet is a mineral in the source, because the size of the heavy REEs (Ho Lu) are more compatible with crystallographic sites in garnet. Using equations that describe how trace elements are partitioned by solids and liquids, concentrations of REEs in melts from garnet peridotite can be calculated. These are shown in the diagram, where REE concentrations have been normalized by dividing the concentration of each element by the concentration found in chondritic meteorites. This produces a REE pattern. Note that the low % melts have Light REE enriched patterns, because the low atomic weight REEs (La - Eu) are enriched over the heavier REEs. Next, we plot the ratio of a highly incompatible element, like La, to a less incompatible element, like Sm, versus the concentration of the highly incompatible element. In the case shown, La/Sm ratio versus La concentration for each % melting. Note the steep slope of the curves connecting the points. As we'll see in our discussion of crystal fractionation, the ratios of incompatible elements do not change much with crystal fractionation, and therefore produce a trend with a less steep slope. This gives us a method for distinguishing between partial melting and crystal fractionation as the process responsible for magmatic differentiation.

53 Crystal Fractionation In our discussion of phase diagrams we saw how liquid compositions can change as a result of removing crystals from the liquid as they form. In all cases except a eutectic composition, crystallization results in a change in the composition of the liquid, and if the crystals are removed by some process, then different magma compositions can be generated from the initial parent liquid. If minerals that later react to form a new mineral or solid solution minerals are removed, then crystal fractionation can produce liquid compositions that would not otherwise have been attained by normal crystallization of the parent liquid. Bowen's Reaction Series Norman L. Bowen, an experimental petrologist in the early 1900s, realized this from his determinations of simple 2- and 3-component phase diagrams, and proposed that if an initial basaltic magma had crystals removed before they could react with the liquid, that the common suite of rocks from basalt to rhyolite could be produced. This is summarized as Bowen's Reaction Series. Bowen suggested that the common minerals that crystallize from magmas could be divided into a continuous reaction series and a discontinuous reaction series. The continuous reaction series is composed of the plagioclase feldspar solid solution series. A basaltic magma would initially crystallize a Ca- rich plagioclase and upon cooling continually react with the liquid to produce more Na-rich plagioclase. If the early forming plagioclase were removed, then liquid compositions could eventually evolve to those that would crystallize a Na-rich plagioclase, such as a rhyolite liquid.

54 The discontinuous reaction series consists of minerals that upon cooling eventually react with the liquid to produce a new phase. Thus, as we have seen, crystallization of olivine from a basaltic liquid would eventually reach a point where olivine would react with the liquid to produce orthopyroxene. Bowen postulated that with further cooling pyroxene would react with the liquid, which by this time had become more enriched in H2O, to produce hornblende. The hornblende would eventually react with the liquid to produce biotite. If the earlier crystallizing phases are removed before the reaction can take place, then increasingly more siliceous liquids would be produced. This generalized idea is consistent with the temperatures observed in magmas and with the mineral assemblages we find in the various rocks. We would expect that with increasing SiO2 oxides like MgO, and CaO should decrease with higher degrees of crystal fractionation because they enter early crystallizing phases, like olivines and pyroxenes. Oxides like H2O, K2O and Na2O should increase with increasing crystal fractionation because they do not enter early crystallizing phases. Furthermore, we would expect incompatible trace element concentrations to increase with fractionation, and compatible trace element concentrations to decrease. This is generally what is observed in igneous rock suites. Because of this, and the fact that crystal fractionation is easy to envision and somewhat easy to test, crystal fraction is often implicitly assumed to be the dominant process of magmatic differentiation. Graphical Representation of Crystal Fractionation The effects on chemical change of magma (rock) compositions that would be expected from crystal fractionation can be seen by examining some simple variation diagrams. In a simple case imagine that we have two rocks, A and B, with their SiO2 and MgO concentrations as shown in the diagram. Also plotted is the analysis of olivine contained in rock A. Removal of olivine from Rock A would drive the liquid composition in a straight line away from A. (This is the same idea we used in phase diagrams). If rock B were produced from rock A by fractionation of olivine, then the composition of rock B should lie on the same line. This should also be true of all other variation diagrams plotting other oxides against SiO2.Just like in phase diagrams we can apply the lever rule to determine how much of the olivine had to fractionate from a magma with composition A to produce rock B: %Olivine Fractionated = [y/(x + y)]*100 If olivine fractionation were the process responsible for the change from magma A to magma

55 B, then these proportions would have to be the same on all other variation diagrams as well. In a more complicated case, we next look at what happens if two phases of different composition were involved in the fractionation. Again the rules we apply are the same rules we used in phase diagrams. In this case, a mixture of 50% olivine and 50% pyroxene has been removed from magma C to produce magma D. Note that the liquid composition has to change along a line away from the composition of the mixture of solid phases, through the composition of the original liquid (magma C). Again the lever rule would tell us that the percentage of solids fractionated would be: %solids fractionated = [z/(w + z)]*100 This works well for small steps in the fractionation sequence. In the real world we find that many minerals expected to crystallize from a magma are solid solutions whose compositions will change as the liquid evolves and temperature drops. We can see how this would affect things with the following example. In this case we look at what happens if an Mg-Fe solid solution mineral is removed as temperature falls. The initial magma has high MgO and low SiO2. The solid crystallizing from this magma also has high MgO and low SiO2. Taking the fractionation in small increments, the second magma produced by removing the solids from the original magma will have higher SiO2 and lower MgO. But, the second liquid will be crystallizing a solid with lower MgO and higher, SiO2, so it will evolve along a different path. The net result will be that the variation will show a curved trend on a variation diagram. Thus, a generalization we can make is that in natural magmas we expect the variation to be along smooth curved trends since most of the minerals that crystallize from magmas are solid solutions. Note that different minerals fractionated will produce different trends, but they will still be smooth and curved. Another complication arises if there is a change in the combination of minerals that are fractionating.

56 In the example shown a series of magmas are produced along segment 1 by fractionating a combination of solids with low FeO and low SiO2. The last magma produced along segment 1 of the variation diagram has different mineral phases in equilibrium. These phases (probably including a mineral with high FeO, like magnetite) have a much higher FeO concentration. Removal of these phases from this magma causes the trend of variation to make a sharp bend, and further fractionation causes liquids to evolve along segment 2. Thus, sudden changes in the trends on variation diagrams could mean that there has been a change in the mineral assemblage being fractionated. Trace Elements and Crystal Fractionation As we might expect, elements that are excluded from crystals during fractionation should have their concentrations increase in the fractionated magmas. This is true for trace elements as well. The concentration of incompatible trace elements should thus increase with increasing crystal fractionation, and the concentration of compatible trace elements should decrease with fractionation. To see how this works with incompatible trace elements, we'll look at the REEs. The diagram shows how the REEs behave as calculated from theoretical equations for trace element distribution. Note that the REE patterns produced by higher percentages of crystal fractionation show higher concentrations, yet the patterns remain nearly parallel to one another. Thus, a suite of rocks formed as a result of crystal fractionation should show nearly parallel trends of REE patterns. Referring back to our discussion of REEs during partial melting, recall that we said that during crystal fractionation the ratios of incompatible elements show little change, and that we can use this factor to distinguish between crystal fractionation and partial melting. Mechanisms of Crystal Fractionation In order for crystal fractionation to operate their must be a natural mechanism that can remove crystals from the magma or at least separate the crystals so that they can no longer react with the liquid. Several mechanisms could operate in nature.

57 Crystal Settling/Floating - In general, crystals forming from a magma will have different densities than the liquid. o If the crystals have a higher density than the liquid, they will tend to sink or settle to the floor of the magma body. The first layer that settles will still be in contact with the magma, but will later become buried by later settling crystals so that they are effectively removed from the liquid. o If the crystals have a lower density in the magma, they will tend to float or rise upward through the magma. Again the first layer that accumulates at the top of the magma body will initially be in contact with the liquid, but as more crystals float to the top and accumulate, the earlier formed layers will be effectively removed from contact with the liquid. Inward Crystallization - Because a magma body is hot and the country rock which surrounds it is expected to be much cooler, heat will move outward away from the magma. Thus, the walls of the magma body will be coolest, and crystallization would be expected to take place first in this cooler portion of the magma near the walls. The magma would then be expected to crystallize from the walls inward. Just like in the example above, the first layer of crystals precipitated will still be in contact with the liquid, but will eventually become buried by later crystals and effectively be removed from contact with the liquid. Filter pressing - this mechanism has been proposed as a way to separate a liquid from a crystal-liquid mush. In such a situation where there is a high concentration of crystals the liquid could be forced out of the spaces between crystals by some kind of tectonic squeezing that moves the liquid into a fracture or other free space, leaving the crystals behind. It would be kind of like squeezing the water out of a sponge. This mechanism is difficult to envision taking place in nature because (1) unlike a sponge the matrix of crystals is brittle and will not deform easily to squeeze the liquid out, and (2) the fractures required for the liquid to move into are generally formed by extensional forces and the mechanism to get the liquid into the fractures involves compressional forces. Filter pressing is a common method used to separate crystals from liquid in industrial processes, but has not been shown to have occurred in nature. Magma Mixing

58 If two or more magmas with different chemical compositions come in contact with one another beneath the surface of the Earth, then it is possible that they could mix with each other to produce compositions intermediate between the end members. If the compositions of the magmas are greatly different (i.e. basalt and rhyolite), there are several factors that would tend to inhibit mixing. Temperature contrast - basaltic and rhyolitic magmas have very different temperatures. If they come in contact with one another the basaltic magma would tend to cool or even crystallize and the rhyolitic magma would tend to heat up and begin to dissolve any crystals that it had precipitated. Density Contrast- basaltic magmas have densities on the order of 2600 to 2700 kg/m3, whereas rhyolitic magmas have densities of 2300 to 2500 kg/m3. This contrast in density would mean that the lighter rhyolitic magmas would tend to float on the heavier basaltic magma and inhibit mixing. Viscosity Contrast- basaltic magmas and rhyolitic magmas would have very different viscosities. Thus, some kind of vigorous stirring would be necessary to get the magmas to mix. Despite these inhibiting factors, there is evidence in rocks that magmas do sometimes mix. The smaller the difference in chemical composition between two magmas, the smaller will be the contrasts in temperature, density, and viscosity. If magmas of contrasting composition come in contact and begin to mix some kind of stirring mechanism would first be necessary. Such stirring could be provided by convection, with the hotter magma rising through the cooler magma. Evidence for Mixing Mingling of magmas. If, in the initially stages of such mixing, the magma were erupted, then we might expect to find rocks that show a "marble cake" appearance, with dark colored mafic rock intermingled with lighter colored rhyolitic rock. This, however, is mingling of magmas. Note that differences in color are not always due to differences in composition, so even in rocks that show this banding, mingling of magmas may not have occurred. Disequilibrium Mineral Assemblages. If convective stirring progresses beyond the point of mingling, some evidence might still be preserved if the crystals present in one of the magmas does not completely dissolve or react. This might leave disequilibrium mineral assemblages. For example, if a basaltic magma containing Mg-rich olivine mixed with a rhyolite magma containing quartz, and the magma was erupted before the quartz or olivine could be redissolved or made over into another mineral, then we would

59 produce a rock containing mineral that are out of equilibrium. Reverse Zoning in Minerals. When a mineral is placed in an environment different than the one in which it originally formed, it will tend to react to retain equilibrium. Instead of dissolving completely or remaking their entire composition, solid solution minerals may just start precipitating a new composition that is stable in the new chemical environment or at the new temperature. This can result in zoned crystals that show reversals of the zoning trends. For Example: Mg-Fe solid solution minerals normally zone from Mg-rich cores to Fe-rich rims. If a Fe-rich olivine or pyroxene is mixed into a Mg-rich magma that is precipitating Mg-rich olivine or pyroxene, it may precipitate the more Mg-rich composition on the rims of the added crystals. Analyses of such crystals would reveal a reversal in zoning. Similarly, if a Na-rich plagioclase originally crystallizing from a rhyolitic magma were mixed into a basaltic magma precipitating a Ca-rich plagioclase, a Ca-rich rim may be added to the Na-rich plagioclase. Glass Inclusions. Crystal growth from liquids is sometimes not perfect. Sometimes the crystal grows incompletely trapping liquid inside. If that liquid is quenched on the surface and a thin section is cut through the crystal this trapped liquid will be revealed as glass inclusions in the crystal. Since the glass inclusions should represent the composition of the magma that precipitated the crystal, chemical analysis of glass inclusions give us the composition of the liquid in which the crystal formed. The groundmass may also contain glass representing the composition of the liquid in which the crystal resided at eruption. If the composition of glass inclusions is different from glass in the groundmass, and if the groundmass composition is not what is expected from normal crystallization of the

60 minerals present, this provides evidence of magma mixing. Chemical Evidence. If the mixing process proceeds to the point where other evidence is erased, evidence for mixing will still be preserved in the composition of the mixed magmas. On oxide-oxide variation diagrams mixtures will lie along a straight line. Thus if diagrams show a group of rocks that lie along the same straight line, and the proportional distances are the same on all diagrams, one could hypothesize that the chemical variation resulted from magma mixing. Crustal Assimilation/Contamination Because the composition of the crust is generally different from the composition of magmas which must pass through the crust to reach the surface, their is always the possibility that reactions between the crust and the magma could take place. If crustal rocks are picked up, incorporated into the magma, and dissolved to become part of the magma, we say that the crustal rocks have been assimilated by the magma. If the magma absorbs part of the rock through which it passes we say that the magma has become contaminated by the crust. Either of these process would produce a change in the chemical composition of the magma unless the material being added has the same chemical composition as the magma. In a sense, bulk assimilation would produce some of the same effects as mixing, but it is more complicated than mixing because of the heat balance involved. In order to assimilate the country rock enough heat must be provided to first raise the country rock to its solidus temperature where it will begin to melt and then further heat must be added to change from the solid state to the liquid state. The only source of this heat, of course, is the magma itself.

61 Let's imagine a simple case of a pure mineral making up the country rock that is to be assimilated. In order to raise the temperature of the country rock from its initial temperature, Ti, to its melting temperature, Tm, heat must be provided. The amount of heat required is determined by the heat capacity of the rock, Cp (the p subscript stands for constant pressure). dh/dt = Cp or ΔH = Cp (Tm - Ti) Once the temperature has risen to Tm, further heat must be added to melt the rock. This heat is known as the heat of melting, ΔHm, also sometimes called the latent heat of fusion. As stated above the heat required for this process must be supplied by the magma. In our example we'll take a simple case where the magma has a eutectic composition and therefore melts at a single temperature. This time we'll assume that the magma is at a temperature, TL, somewhat above its melting Temperature, Tm. In order to provide the heat for assimilation it would first have to cool. The amount of heat it would give up would then be: ΔH = CpLiquid (Tm - Ti) Once it reached Tm the only other source of heat must be provided by crystallization to release the latent heat of crystallization, ΔHc. Note that in the case shown here, the total heat released by the magma in cooling to Tm and crystallizing is still not enough to melt the country rock. Furthermore, in this process the magma has completely crystallized, so assimilation cannot take place. There are two ways to overcome this problem: (1) If the initial temperature of the magma were much higher, then it could provide the heat by simply cooling to its melting temperature. This is unrealistic, however, because magmas are probably relatively close to or below their liquidus temperature after having passed through cooler country rocks. (2) If the country rock had an initial temperature closer to its melting temperature, less heat would have to be provided by the magma. This could happen if there were successive batches of magma passing through and releasing heat into the country rock. Nevertheless, this heat budget analysis illustrates the difficulty involved in bulk assimilation of

62 country rock by magmas, and makes the process less attractive as a process to explain the chemical diversity of a suite of rocks. Note also that the heat budget will still likely involve crystallizing some of the magma, so if assimilation takes place it will likely involve a combined process of crystal fractionation and assimilation. In a more realistic natural situation things will be slightly different because both the country rock and the magma will melt/crystallize over a range of temperatures, rather than at a single temperature. Even still, the amount of heat required to melt the rock must be provided over a relatively narrow range of temperature. But, partial assimilation of the country rock would be possible because the country rock would only have to be partially melted to produce a liquid that could mix with the magma. In this case we would say that the magma has been contaminated by the country rock. Evidence for Assimilation/Contamination As magma passes upward through the crust pieces of the country rock through which it passes may be broken off and assimilated by the magma. Just as in magma mixing, various stages of this process may be preserved in the magma and rock that results. Xenoliths (meaning foreign rock) are pieces of rock sometimes found as inclusions in other rocks. The presence of xenoliths does not always indicate that assimilation has taken place, but if the xenoliths show evidence of having been disaggregated with their minerals distributed thought the rest of the rock it is likely that some contamination of the magma has taken place. This may result in disequilibrium mineral assemblages and reversely zoned minerals, just as in the case of magma mixing. And, if the assimilation goes to completion, with all of the xenoliths being dissolved in the magma, the only evidence left may be chemical, and again similar to the straight line mixing patterns produced by mixing. Perhaps the best evidence of assimilation/contamination comes from studies of radiogenic isotopes. Here we give an example using the systematics of the Rb - Sr system. 87 Rb is a radioactive isotope that decays to 87Sr with a half life of 47 billion years. Because Rb is an incompatible element, it has been extracted from the mantle by magmas and added to the crust. Thus the concentration of Rb in the crust (avg. about 100 ppm) is much higher than it is in the mantle (avg. about 4 ppm).

63 86 Because 87Rb decays to produce 87Sr and because there is more Rb in the crust than in the mantle, the 87Sr/ 86Sr of the crust has, over time, changed to much higher values than the 87Sr/86Sr ratio in the mantle. The 87Sr/86Sr ratio of the mantle is generally in the range between Thus, rocks derived from melting of the mantle should have 87Sr/86Sr ratios in this range. 87 Sr/86Sr ratios of crustal rocks will depend on their age and concentration of Rb. Older crustal rocks will have high values of 87Sr/86Sr in the range , younger crustal rocks having been recently derived from the mantle will 87Sr/86Sr ratios more similar to the mantle. If mantle derived magmas assimilate or are contaminated by older crustal rocks, then we would expect to find ratios of 87Sr/86Sr in these contaminated rocks that are higher than those found in the mantle and extend up to values found in older crustal rocks. For a suite of rocks affected by contamination, 87Sr/86Sr ratio plotted against Sr concentration would plot along a hyperbolic mixing curve. Note that magma mixing could produce similar trends if the two end member have different concentrations of Sr and 87Sr/86Sr ratios. Crystal fractionation, on the other hand does not change radiogenic isotopic ratios. Sr is a stable, non radiogenic isotope whose concentration does not change with time. Liquid Immiscibility Liquid immiscibility is where liquids do not mix with each other. We are all familiar with this phenomenon in the case of oil and water/vinegar in salad dressing. We have also discussed immiscibility in solids, for example in the alkali feldspar system. Just like in the alkali feldspar system, immiscibility is temperature dependent.

64 For example, in a two component system if there is a field of immiscibility it would appear as in the diagram shown here. Cooling of a liquid with a composition of 25%B & 75%A would eventually result in the liquid separating into two different compositions. With further cooling one liquid would become more enriched in A and the other more enriched in B. Eventually both liquids would reach a temperature where crystals of A would start to form. Note that both liquids would be in equilibrium with crystals of A at the same temperature. Further cooling would result in the disappearance of the A-rich liquid. This points out two important properties of immiscible liquids. 1. If immiscible liquids are in equilibrium with solids, both liquids must be in equilibrium with the same solid compositions. 2. Extreme compositions of the two the liquids will exist at the same temperature. Liquid immiscibility was once thought to be a mechanism to explain all magmatic differentiation. If so, requirement 2, above, would require that siliceous liquids and mafic liquids should form at the same temperature. Since basaltic magmas are generally much hotter than rhyolitic magmas, liquid immiscibility is not looked upon favorably as an explanation for wide diversity of magmatic compositions. Still, liquid immiscibility is observed in experiments conducted on simple rock systems. For example, in the system Fo-An-Qz a field of immiscible liquids is observed for compositions rich in SiO2. But these compositions are outside of the range of compositions that occur in nature. This is true of almost all simple systems wherein liquid immiscibility has been observed. There are however, three exceptions where liquid immiscibility may play a role. 1. Sulfide liquids may separate from mafic silicate magmas.

65 2. Highly alkaline magmas rich in CO2 may separate into two liquids, one rich in carbonate, and the other rich in silica and alkalies. This process may be responsible for forming the rare carbonatite magmas 3. Very Fe-rich basaltic magmas may form two separate liquids - one felsic and rich in SiO2, and the other mafic and rich in FeO. Combined Processes As pointed out previously, if any of these process are possible, then a combination of the process could act to produce chemical change in magmas. Thus, although crystal fractionation seems to be the dominant process affecting magmatic differentiation, it may not be the only processes. As we have seen, assimilation is likely to accompanied by crystallization of magmas in order to provide the heat necessary for assimilation. If this occurs then a combination of crystal fraction and assimilation could occur. Similarly, magmas could mix and crystallize at the same time resulting in a combination of magma mixing and crystal fractionation. In nature, things could be quite complicated. Magmatic Differentiation The different types of magmas produced in the different tectonic regimes, are dependent on three factors: source rock composition, melting conditions and processes, and crystallization conditions and processes. The composition of the source rock, the degree or extent to which the source rock is melted, the pressure and and temperature conditions under which melting takes place, and the processes by which melting and separation of the melt take place may affect the composition of the melt or magma produced. Many of these same variables will also effect crystallization. The conditions and processes of melting and crystallization may change over time. Due to the large number of variables, a wide range of igneous rocks compositions and textures are possible. As a magma crystallizes, the magma becomes depleted in the elements that are entering the crystallizing minerals and so the melt changes composition over time. Magmatic differentiation is the process whereby one parent magma composition can produce a number of igneous rock compositions through mineral crystallization. As a cooling melt changes composition, the minerals that are in equilibrium with it (i.e. that are stable in the melt at the temperature and pressure conditions of crystallization) typically either change composition and/or change to structurally-more complex minerals as in Bowen's Reaction Series. Magmatic differentiation occurs through equilibrium crystallization or fractional crystallization or most commonly some crystallization scheme intermediate between the two end member processes.

66 Equilibrium crystallization is the process whereby every crystallizing mineral has an opportunity to react with the residual melt and change composition or mineral structure to remain in equilibrium with the melt. Fractional crystallization is the process whereby crystals once produced are instantly isolated from the melt and prevented from equilibrating with the liquid from which they crystallized, resulting in a series of residual liquids of more extreme compositions than would have resulted from equilibrium crystallization. In reality, the crystals in most crystallizing magmas are neither perfectly in equilibrium with the residual melt or perfectly isolated from reaction with the melt immediately after crystallization. Magmas also change composition as a result of assimilation, magma mixing, or separating into immiscible liquids. These processes all occur as a result of diffusion of elements or between the magma and other materials with different compositions. Assimilation is a process whereby magma composition changes as a result of the incorporation and digestion of wall or country rock in the magma. If a high temperature magma intrudes a shale (an aluminous clay mineral-rich sedimentary rock), partial or total melting of the shale in contact with the magma, may occur. The aluminum content of the magma would increase (the aluminum content in the shale far exceeds the aluminum content in the original magma composition) and aluminum-rich minerals such as garnet, corundum, cordierite, and muscovite may occur. Magma mixing occurs when two or more magmas of different compositions mix, or partially mix to produce a magma of a different composition. Magmatic differentiation may produce a melt that separates into two or more immiscible fluids. The most common types of immiscible geologic melts are sulfide-rich and silicate-rich melts. The first geologic evidence for the existence of immiscible fluids was discovered in lunar rocks!

67 (COMPILED BY GDC HANDWARA)

68 Lecture notes,2nd semester,unit -2 ORIGIN OF MAGMAS Magmas do not form everywhere beneath the surface of the Earth. This is evident from looking at the world distribution of volcanoes. Thus, magmas must require special circumstances in order to form. Before we talk about how and where magmas form, we first look at the interior structure of the Earth. The Earth's Internal Structure Evidence from seismology tells us that the Earth has a layered structure. Seismic waves generated by earthquakes travel through the Earth with velocities that depend on the type of wave and the physical properties of the material through which the waves travel. Types of Seismic Waves Body Waves - travel in all directions through the body of the Earth. There are two types of body waves: o P - waves - are Primary waves. They travel with a velocity that depends on the elastic properties of the rock through which they travel. Vp = [(K + 4/3μ)/ρ Where, Vp is the velocity of the P-wave, K is the incompressibility of the material, μ is the rigidity of the material, and ρ is the density of the material. P-waves are the same thing as sound waves. They move through the material by compressing it, but after it has been compressed it expands, so that the wave moves by compressing and expanding the material as it travels. Thus the velocity of the P-wave depends on how easily the material can be compressed (the incompressibility), how rigid the material is (the rigidity), and the density of the material. P-waves have the highest velocity of all seismic waves and thus will reach all seismographs first. o S-Waves - Secondary waves, also called shear waves, travel with a velocity that depends only on the rigidity and density of the material through which they travel: Vp = [(μ)/ρ] S-waves travel through material by shearing it or changing its shape in the direction perpendicular to the direction of travel. The resistance to shearing of a material is the property called the rigidity. It is notable that liquids have no rigidity, so that the velocity of an S-wave is zero in a liquid. (This point will become important later). Note that Swaves travel slower than P-waves, so they will reach a seismograph after the P-wave.

69 Surface Waves - Surface waves differ from body waves in that they do not travel through the Earth, but instead travel along paths nearly parallel to the surface of the Earth. Surface waves behave like S-waves in that they cause up and down and side to side movement as they pass, but they travel slower than S-waves and do not travel through the body of the Earth. Thus they can give us information about the properties of rocks near the surface, but not about the properties of the Earth deep in the interior. Because seismic waves reflect from and refract through boundaries where there is sudden change in the physical properties of the rock, by tracing the waves we can see different layers in the Earth. This allows us to look at the structure of the Earth based on layers of differing physical properties. Note that we know that density must increase with depth in the Earth because the density of crustal rocks are about 2,700 kg/m3 and the average density of the Earth is about 5,200 kg/m3. Also note from the velocity equations that if density increases, wave velocity decreases. Thus, the other properties, incompressibility and rigidity must increase with depth in the Earth at a greater rate than density increases. Once we know the seismic wave velocities throughout the Earth, then we can perform experiments on different possible materials and make estimates of what the chemical composition. Thus, we can also divide the Earth into layers of differing chemical composition. Layers of Differing Chemical Composition Crust - variable thickness and composition

70 o Continental km thick, underlies all continental areas, has an average composition that is andesitic. o Oceanic 8-10 km thick, underlies all ocean basins, has an average composition that is basaltic Mantle km thick, made up of a rock called peridotite (Olivine + Opx + Cpx). Evidence comes from Seismic wave velocities, experiments, and peridotite xenoliths (foreign rocks) brought to the surface by magmas. Experimental evidence suggests that the mineralogy of peridotite changes with depth (ant thus pressure) in the Earth. At low pressure, the mineral assemblage is Olivine + Cpx + Opx + Plagioclase (plagioclase peridotite). At higher pressure the assemblage changes to Olivine + Cpx + Opx + Spinel [(Mg,Fe+2) (Cr, Al, Fe+3)2O4] (spinel peridotite). At pressures above about 30 kilobars, the assemblage changes to Olivine + Cpx + Opx + garnet (garnet peridotite). This occurs because Al changes its coordination with increasing pressure, and thus new minerals must form to accommodate the Al. At greater depths, such as the 400 km discontinuity and the 670 km discontinuity, olivine and pyroxene likely change to high pressure polymorphs. Despite these changes in mineral assemblage, the chemical composition of the mantle does not appear to change much in terms of its major element composition. Core km radius, made up of Iron (Fe) and small amount of Nickel (Ni). Evidence comes from seismic wave velocities, experiments, and the composition of iron meteorites, thought to be remnants of other differentiated planets that were broken apart due to collisions.

71 Layers of Differing Physical Properties Lithosphere - about 100 km thick (up to 200 km thick beneath continents, thinner beneath oceanic ridges and rift valleys), very brittle, easily fractures at low temperature. Note that the lithosphere is comprised of both crust and part of the upper mantle. The plates that we talk about in plate tectonics are made up of the lithosphere, and appear to float on the underlying asthenosphere. Asthenosphere - about 250 km thick - solid rock, but soft and flows easily (ductile). The top of the asthenosphere is called the Low Velocity Zone (LVZ) because the velocities of both P- and S-waves are lower than the in the lithosphere above. But, not that neither P- nor S-wave velocities go to zero, so the LVZ is not completely liquid. Mesosphere - about 2500 km thick, solid rock, but still capable of flowing. Outer Core km thick - liquid. We know this because S-wave velocities are zero in the outer core. If Vs = 0, this implies μ = 0, and this implies that the material is in a liquid state. Inner core km radius, solid Where do Magmas Come From? Magmas are not likely to come from the only part of the Earth that is in a liquid state, the outer core, because it does not have the right chemical composition. The outer core is made mostly of Fe with some Ni, magmas are silicate liquids. In the ocean basins, magmas are not likely to come from melting of the oceanic crust, since most magmas erupted in the ocean basins are basaltic. To produce basaltic magmas by melting of the basaltic oceanic crust would require nearly 100% melting, which is not likely. In the continents, both basaltic and rhyolitic magmas are erupted and intruded. Basaltic magmas are not likely to have come from the continental crust, since the average composition is more siliceous, but more siliceous magmas (andesitic - rhyolitic) could come from melting of the continental crust. Basaltic magmas must come from the underlying mantle. Thus, with the exception of the continents, magmas are most likely to originate in the mantle from melting of mantle peridotite. Origin of Magmas Again, magmas do not form everywhere beneath the surface, so special circumstances are necessary.

72 Temperature varies with depth or pressure in the Earth along the geothermal gradient. The normal geothermal gradient is somewhat higher beneath the oceans than beneath the continents, at least at shallow levels. If we compare the normal geothermal gradients with the experimentally determined phase diagram for peridotite containing little water or carbon dioxide, we find that the peridotite solidus temperature is everywhere higher than the normal geothermal gradients. Thus, under normal conditions the mantle is solid, as we would suspect from the seismic evidence. Thus, in order to generate a melt, either we must find a way to increase the geothermal gradient so that it is above the peridotite solidus or reduce the temperature of the peridotite solidus. In either case note that all we have to do is get the temperature in some part of the Earth, as expressed by the geothermal gradient, into the field of partial melt. Partial melting is the most likely case because it requires less of an increase in temperature or less of a decrease in the peridotite solidus. Once a partial melt has formed, the liquid portion can be easily separated from the remaining solids since liquids are more mobile and, in general, have a lower density than solids. Raising the Geothermal Gradient Radioactive Heat - Elements like U, Th, K, and Rb have radioactive isotopes. During radioactive decay, sub-atomic particles are released by the decaying isotope and move outward until they collide with other atomic particles. Upon collision, the kinetic energy of the moving particles is converted to heat. If this heat cannot be conducted away, then the temperature will rise. Most of the heat within the Earth is generated by radioactive decay, and this is the general reason why temperature increases with depth in the Earth. But most the radioactive isotopes are concentrated in the crust. Although there are areas in the continental crust where high concentrations of radioactive elements have locally raised the temperature, at least high enough to cause metamorphism, this is a rare occurrence. It is even more unlikely that areas of high concentration develop within the mantle. Thus, concentrations of radioactive elements is not likely to cause melting.

73 Frictional Heat - In areas where rocks slide past one another, such as at the base of the lithosphere, on at subduction zones, heat could be generated by friction. If this heat cannot be conducted away fast enough, then it may cause a localized rise in temperature within the zone where the sliding or shearing is taking place. This could cause a localized spike on the geothermal gradient that could cause local temperatures to rise above the solidus. Decompression due to Convection - Convection is a form of heat transfer wherein the heat moves with the material. Convection can be induced if the temperature gradient is high enough that material at depth expands so that its density is lower than the material above it. This is an unstable situation and the hotter, lower density material will rise to be replaced by descending cooler material in a convection cell. The rate of convection depends on both on the temperature gradient and the viscosity of the material (note that solids convect, but the rate is lower than in liquids because solids have higher viscosity). In the Earth, temperature gradients appear to be high enough and viscosity low enough for convection to occur. Plate tectonics appears to be driven by convection in some form.

74 Anywhere there is a rising convection current, hotter material at depth will rise carrying its heat with it. As it rises to lower pressure (decompression) it will cool somewhat, but will still have a temperature higher than its surroundings. Thus, decompression will result in raising the local geothermal gradient. If this new geothermal gradient reaches temperatures greater than the peridotite solidus, partial melting and the generation of magma can occur. This mechanism is referred to as decompression melting. Lowering the Solidus Temperature As we saw in our discussion of phase diagrams, mixtures of components begin melting at a lower temperature than the pure components. In a two component system addition of a third component reduces both the solidus and liquidus temperatures. This suggests that if something can be added to the mantle, it could cause the solidus and liquidus temperatures to be lowered to the extent that the solidus could become lower than the geothermal gradient and result in partial melting, without having to raise the geothermal gradient. Such a melting mechanism is referred to as flux melting. It's difficult to imagine how solid components could be added to the mantle. But volatile components, for example H2O and CO2, because of their high mobility, could be added to the mantle, particularly at subduction zones. Oceanic crust is in contact with sea water, thus water could be in oceanic crust both due to weathering, which produces hydrous minerals like clay minerals, and could be in the pore spaces in the rock. Oceanic sediments eventually cover the basaltic oceanic crust produced at oceanic ridges. Much of this sediment consists of clay minerals (which contain water) and carbonate minerals (which contain carbon dioxide). As the oceanic lithosphere descends into the mantle at a subduction zone, it will be taken to increasingly higher temperatures as it gets deeper. This will result in metamorphism of both the basalt and the sediment. As we will see later in our discussion of metamorphism, metamorphism is essentially a series of dehydration and decarbonation reactions, i.e. chemical reactions that transform hydrous and carbonate minerals into nonhydrous minerals and give up H2O and CO2 as a fluid phase.

75 Addition of this fluid phase, either to the subducted lithosphere or the mantle overlying the subducted lithosphere could lower the solidus and liquidus temperatures enough to cause partial melting. Crustal Anatexis In the continental crust, it is not expected that the normal geothermal gradient will be high enough to cause melting despite the fact that hydrous and carbonate minerals occur in many continental rocks. Furthermore, because continental rocks are at low temperature and have a very high viscosity, convective decompression is not likely to occur. Yet, as we will see, there is evidence that continental crustal rocks sometimes melt. This is called crustal anatexis. The following scenario is one mechanism by which crustal anatexis could occur. Basaltic magmas, generated in the mantle, by flux melting, decompression melting or frictional heat, rise into the crust, carrying heat with them. Because basaltic liquids have a higher density than crust, they may not make it all the way to the surface, but instead intrude and cool slowly at depth. Upon cooling the basaltic magmas release heat into the crust, raising the geothermal gradient (increasing the local temperature). Successive intrusions of mantle-derived mantle into the same area of the crust may cause further increases in temperature, and eventually cause the geothermal gradient to become higher than the wet solidus of the crustal material, resulting in a partial melt of the crust.

76 Magmatism and Plate Tectonics From the discussion above it should be obvious that magmatism is closely related to plate tectonics. The diagram below summarizes melting mechanisms that occur as a result of plate tectonics and may be responsible for the generation of magmas in a variety of plate tectonic settings, such as oceanic ridges, near subduction zones, and at rift valleys. Diverging Plate Boundaries Diverging plate boundaries are where plates move away from each other. These include oceanic ridges or spreading centers, and rift valleys. Oceanic Ridges are areas where mantle appears to ascend due to rising convection currents. Decompression melting could result, generating magmas that intrude and erupt at the oceanic ridges to create new oceanic crust. Iceland is one of the few areas where the resulting magmatism has been voluminous

77 enough to built the oceanic ridge above sea level. Continental Rift Valleys or Extensional Zones are areas, usually located in continental crust where extensional deformation is occurring. These areas may be incipient spreading centers and may eventually evolve into oceanic ridges, such as has occurred in the Red Sea region. Whether or not they develop into spreading centers, they are likely caused by mantle upwelling below the zone of extension. Mantle upwelling may result in decompression melting of the mantle, and could induce crustal anatexis. A good example of a continental rift valley is the East African Rift Valley. Another example is the Rio Grande Rift in Colorado and New Mexico, which is part of a larger region of extension that includes much of the western U.S. and is called the Basin and Range Province (Eastern California & Oregon, Nevada, Utah, Arizona, & New

78 Mexico). Converging Plate Boundaries Converging plate boundaries are where plates run into each other. The most common type are where oceanic lithosphere subducts. Several mechanisms could contribute to the generation of magmas in this environment (see diagram at top of this section). 1. Frictional heating is likely to occur along the boundary between the subducted plate and the overlying mantle wedge. 2. Flux melting of either the subducted lithosphere or the overlying mantle wedge could occur as a result of the release of volatiles as the subducted plate heats and metamorphoses producing water and/or carbon dioxide fluids. 3. The process of subduction may drag the overlying mantle wedge down with it. In order to replace the mantle dragged down in this process, part of the mantle wedge will have to rise. This upwelling of the mantle could result in decompression melting. If an oceanic lithospheric plate subducts beneath another oceanic lithospheric plate, we find island arcs on the surface above the subduction zone. If an oceanic plate subducts beneath a plate composed of continental lithosphere, we find continental margin arcs. If magma generated near the subduction zone intrudes and cools in the crust, it could induce crustal anatexis.

79 In areas where two continental lithospheric plates converge fold-thrust mountain ranges develop as the result of compression. If water-bearing crustal rocks are pushed to deeper levels where temperatures are higher, crustal anatexis may result. Intraplate Magmatism & Hot Spots There are a few areas where magmatism does not appear to be related to converging or diverging plate boundaries. These areas occur in the middle of plates, usually far from the plate boundaries. This phenomenon is referred to as intraplate magmatism. Intraplate magmatism is thought to be caused by hot spots formed when thin plumes of mantle material rise along narrow zones from deep within the mantle. The hot spot remains stationary in the mantle while the plate moves over the hot spot. Decompression melting caused by the upwelling plume produces magmas that form a volcano on the sea floor above the hot spot. The volcano remains active while it is over the vicinity of the hot spot, but eventually plate motion results in the volcano moving away from the plume and the volcano becomes extinct and begins to erode.

80 Because the Pacific Plate is one of the faster moving plates, this type of volcanism produces linear chains of islands and seamounts, such as the Hawaiian - Emperor chain, the Line Islands, the Marshall-Ellice Islands, and the Austral seamount chain. (PREPARED BY GDC HANDWARA)