Odale-Articles-Introduction

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IMPACT CRATER EXPLORATIONS

by: Charles O'Dale

Introduction

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Recently I have had the privilege of presenting to the Ottawa RASC images of the many terrestrial meteorite craters that I have explored from the air in my Cessna 177B Cardinal. To accomplish this I have combined three of my many hobbies (astronomy, geology and aviation) into one great hobby, astro-geology. This is the relatively new science for the study of terrestrial impact craters. I am in the process of documenting all of North America’s meteorite craters in this way. My aerial explorations have given me an appreciation of how asteroid impacts have combined with tectonics and natural erosion to shape the world as we see today.

This is my trusty starship, C-GOZM, pictured here over the Arecibo Radio telescope in Puerto Rico. We were on our way to Antigua see the 1998 total solar eclipse. I like to say that I flew all the way to Puerto Rico to explore a crater only to find that someone had built a radio telescope in it!

Crater Classification

In this website, I will present the craters that I have explored in the order of their size, starting with simple craters, moving to complex craters and finishing with multiring basins. The transition size between simple to complex craters is 2km in sediments and 4km in crystalline rocks. The transition size between complex to ringed basin craters is 10 to 50 km.
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The central peak of the complex crater is formed as a result of uplift of material stratigraphically beneath the crater, which rebounds in response to compression caused by the impact. From Melosh (1989)
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Peak ring craters are the more complex form and occur as larger (greater than 50km) complex craters. From Melosh (1989)

References

For some of the craters I will briefly delve into the basics of their formation but it is beyond the scope of this article to present the complete physics of impact crater formation. For this I will recommend the following references:



Shock Metamorphism

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I refer to the shock metamorphism that is noted at each crater. This is the extreme physical conditions that are imposed by intense shock waves on the impacted rocks. Such shock waves are produced naturally only by the hypervelocity impacts of extraterrestrial objects and cause large volumes of target rock to be shattered, deformed, melted and even vaporized in a few seconds. Most impact rocks are easy to link to their source material from their chemical, mineralogical and isotopic characteristics. Impactor shock pressure and metamorphism in nonporous crystalline target rock is illustrated in the diagram and listed with the noted increasing impact pressure (from Traces of Catastrophe). The standard unit of pressure is the pascal, abbreviated Pa, which is equivalent to 1 kg/m2 (1 kilogram per square meter). A GPa is a gigapascal (giga means billion), a measurement of pressure, and is equal to 10,000 times the pressure at the Earth's surface. The pressure at the center of the Earth is about 361 GPa, or more than four million times surface pressure.


Fracturing and Brecciation

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1 GPa to 5 GPa - Fracturing (left) and Brecciation (right) - Generally, impact breccias are made up of fragments of the target rocks, containing various ratios of impact melt and shocked mineral inclusions.



Shatter Cones

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2 GPa to 30 GPa - Shatter cones are shock-deformation features that form from impact pressures of typically 2-10 GPa up to ~30 GPa. They represent the only distinctive and unique shock-deformation feature that develops on a megascopic scale (e.g., hand sample to outcrop scale). They appear in outcrops as distinctively curved striated fractures that typically form partial or complete conical structures (image). They are commonly found beneath impact crater floors, usually in the central uplifts of complex impact structures, but they may also be observed in isolated rock fragments within brecciated units.


Comminution and Fracturing

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<10 GPa - The shocked Coconino Sandstone (top) is weakly shocked sandstone that lacks remnant porosity and contains abundant grain comminution and fracturing. Note the "rock flour" on the shocked sample. The unshocked Coconino Sandstone (bottom) consists of a fine to medium-grained, moderately well-sorted, rounded quartz arenite with ~ 20 vol% porosity. The shocked sample is from the Barringer impact structure. Coconino sandstone layers are typically buff to white in color. It consists primarily of sand deposited by eolian processes (wind-deposited) approximately 260 million years ago.

Microscopic planar deformation features

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~8 GPa to 25 GPa - Quartz is a mineral which retains particularly well the memory of the extreme pressures induced during impact. Shocked quartz grains contain very fine lamellae of amorphous silica. Experiments have concluded that such lamellae can only be produced by a rapid pressure change of > 10,000 atmospheres.

Planar deformation features are impact diagnostic and cannot form in any other geological environment.

Mineral Transformation

25 GPa to 40 GPa - Transformation of individual minerals to amorphous phases without melting. Graphite can be converted to diamond. Quartz can be converted to stishovite at shock pressures of >12-15 GPa and to coesite at >30 GPa. The identification of coesite and stishovite at several sites in the early 1960s provided one of the earliest criteria for establishing the impact origin of several structures. Quartz and feldspars are the most common examples of minerals converted to diaplectic glasses by shock waves. Diaplectic plagioclase feldspar glass, called maskelynite, was in fact observed in meteorites more than a century before it was discovered in shocked terrestrial rocks.

Partial Melting

35 GPa to 60 GPa - Selective partial melting of individual minerals.

Melt

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60 GPa to 100 GPa - Complete melting of all minerals - these examples are from the Sudbury impact structure.



Vaporization

100 GPa - Complete rock vaporization.

Pseudotachylite

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Pseudotachylite is a type of impactite breccia that occurs in target rocks of large impact structures. It is defined as a breccia having the aspect and the black color of a volcanic rock (a tachylite), and that is formed when a high pressure (caused by an earthquake or a meteorite impact) is applied to a rock and abruptly released, which cause the rock to partly melt. Chemical studies of pseudotachylites have shown that they correspond closely to the adjacent host rocks, indicating that they have formed essentially in place by locally generated cataclastic milling and/or frictional melting processes. Bodies of techtonic pseudotachylite tend to be linear and less than a few metres wide. Impact-produced pseudotachylites form more irregular bodies, some of which may reach tens to hundreds of metres in size. This example of pseudotachylite is from the Sudbury impact structure.



Suevite

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Suevite is a melt-fragment breccia composed of discrete fragments of rocks and minerals, together with bodies of melt, in a clastic matrix of similar but finer-grained materials. Many of the rock and mineral fragments are highly shocked, and these breccias often provide the most distinctive evidence for a meteorite impact origin of the structures in which they are found. These suevite examples are from the Lake Wanapitei crater.





Radiometric Dating Of Meteorite Craters

I refer to the following dating methods that were used to determine the ages of the various impacts.

Potassium-Argon

The potassium-argon (K-Ar) ratio is “reset” to zero by impacts when the argon gas generated by the decay of potassium is diffused out of material heated by shock. The K-Ar age will then date the impact which affected the target bedrock.

Potassium is an abundant element in the Earth's crust. One isotope, potassium-40, is radioactive and decays to two different daughter products, calcium-40 and argon-40, by two different decay methods. The production ratio of these two daughter products is precisely known, and is always constant: 11.2% becomes argon-40 and 88.8% becomes calcium-40. It is possible to date some rocks by the potassium-calcium method, but this is not often done because it is hard to determine how much calcium was initially present. Argon, on the other hand, is a gas. Whenever rock is melted to become impact melt, magma or lava, the argon tends to escape. Once the molten material hardens, it begins to trap the new argon produced since the hardening took place. In this way the potassium-argon clock is clearly reset when an igneous rock is formed.

In its simplest form, the geologist simply needs to measure the relative amounts of potassium-40 and argon-40 to date the rock. The age is given by a relatively simple equation:

t = h x ln[1 + (argon-40)/(0.112 x (potassium-40))]/ln(2)

where t is the time in years, h is the half-life, also in years, and ln is the natural logarithm.

However, in reality there is often a small amount of argon remaining in a rock when it hardens. This is usually trapped in the form of very tiny air bubbles in the rock. One percent of the air we breathe is argon. Any extra argon from air bubbles may need to be taken into account if it is significant relative to the amount of radiogenic argon (that is, argon produced by radioactive decays). This would most likely be the case in either young rocks that have not had time to produce much radiogenic argon, or in rocks that are low in the parent potassium. One must have a way to determine how much air-argon is in the rock. This is rather easily done because air-argon has a couple of other isotopes, the most abundant of which is argon-36. The ratio of argon-40 to argon-36 in air is well known, at 295. Thus, if one measures argon-36 as well as argon-40, one can calculate and subtract off the air-argon-40 to get an accurate age.

One of the best ways of showing that an age-date is correct is to confirm it with one or more different dating method(s). Although potassium-argon is one of the simplest dating methods, there are still some cases where it does not agree with other methods. When this does happen, it is usually because the gas within bubbles in the rock is from deep underground rather than from the air. This gas can have a higher concentration of argon-40 escaping from the melting of older rocks. This is called parentless argon-40 because its parent potassium is not in the rock being dated, and is also not from the air. In these slightly unusual cases, the date given by the normal potassium-argon method is too old. However, scientists in the mid-1960s came up with a way around this problem, the argon-argon method.

Argon-Argon

This method uses exactly the same parent and daughter isotopes as the potassium-argon method. In effect, it is a different way of telling time from the same clock. Instead of simply comparing the total potassium with the non-air argon in the rock, this method has a way of telling exactly what and how much argon is directly related to the potassium in the rock.

In the argon-argon method the rock is placed near the center of a nuclear reactor for a period of hours. A nuclear reactor emits a very large number of neutrons, which are capable of changing a small amount of the potassium-39 into argon-39. Argon-39 is not found in nature because it has only a 269-year half-life. (This half-life doesn't affect the argon-argon dating method as long as the measurements are made within about five years of the neutron dose). The rock is then heated in a furnace to release both the argon-40 and the argon-39 (representing the potassium) for analysis. The heating is done at incrementally higher temperatures and at each step the ratio of argon-40 to argon-39 is measured. If the argon-40 is from decay of potassium within the rock, it will come out at the same temperatures as the potassium-derived argon-39 and in a constant proportion. On the other hand, if there is some excess argon-40 in the rock it will cause a different ratio of argon-40 to argon-39 for some or many of the heating steps, so the different heating steps will not agree with each other.

Rubidium-Strontium

In nearly all of the dating methods, except potassium-argon and the associated argon-argon method, there is always some amount of the daughter product already in the rock when it cools. The Rubidium-Strontium dating method reveals how much of the daughter product was already in the rock when it cooled and hardened.

The nuclide rubidium-87 decays, with a half life of 48.8 billion years, to strontium-87. Strontium-87 is a stable element; it does not undergo further radioactive decay. (Do not confuse with the highly radioactive isotope, strontium-90.) Strontium occurs naturally as a mixture of several nuclides, including the stable isotope strontium-86. If three different strontium-containing minerals form at the same time in the same magma, each strontium containing mineral will have the same ratios of the different strontium nuclides, since all strontium nuclides chemically behave the same. (Note that this does not mean that the ratios are the same everywhere on earth. It merely means that the ratios are the same in the particular magma from which the test sample was later taken.) As strontium-87 forms, its ratio to strontium-86 will increase. Strontium-86 is a stable element that does not undergo radioactive change. In addition, it is not formed as the result of a radioactive decay process. The amount of strontium-86 in a given mineral sample will not change. Therefore the relative amounts of rubidium-87 and strontium-87 can be determined by expressing their ratios to strontium-86: Rb-87/Sr-86 and Sr87/Sr-86 We measure the amounts of rubidium-87 and strontium-87 as ratios to an unchanging content of strontium-86.

Uranium-Lead

The uranium-lead method is the longest-used dating method. It was first used in 1907, about a century ago. The uranium-lead system is more complicated than other parent-daughter systems; it is actually several dating methods put together. Natural uranium consists primarily of two isotopes, U-235 and U-238, and these isotopes decay with different half-lives to produce lead-207 and lead-206, respectively. In addition, lead-208 is produced by thorium-232. Only one isotope of lead, lead-204, is not radiogenic. The uranium-lead system has an interesting complication: none of the lead isotopes is produced directly from the uranium and thorium. Each decays through a series of relatively short-lived radioactive elements that each decay to a lighter element, finally ending up at lead. Since these half-lives are so short compared to U-238, U-235, and thorium-232, they generally do not affect the overall dating scheme. The result is that one can obtain three independent estimates of the age of a rock by measuring the lead isotopes and their parent isotopes.

The uranium-lead system in its simpler forms, using U-238, U-235, and thorium-232, has proved to be less reliable than many of the other dating systems. This is because both uranium and lead are less easily retained in many of the minerals in which they are found. Yet the fact that there are three dating systems all in one allows scientists to easily determine whether the system has been disturbed or not. Using slightly more complicated mathematics, different combinations of the lead isotopes and parent isotopes can be plotted in such a way as to minimize the effects of lead loss. One of these techniques is called the lead-lead technique because it determines the ages from the lead isotopes alone. Some of these techniques allow scientists to chart at what points in time metamorphic heating events have occurred, which is also of significant interest to geologists.

Dating Methods Used For The Recent 100,000 Years

Thermoluminescence (TL) Dating

A method of dating minerals and pottery. Rather than relying on a half-life, this method relies instead on the total amount of radiation experienced by the mineral since the time it was formed. This radiation causes disorder in the crystals, resulting in electrons dwelling in higher orbits than they originally did. When the sample is heated in the laboratory in the presence of a sensitive light detector, these electrons return to their original orbits, emitting light and allowing an age to be determined by comparison of the amount of light to the radioactivity rate experienced by the mineral. Variations on this method include optically-stimulated luminescence (OSL) and infrared-stimulated luminescence (IRSL) dating.

Chlorine-36 Dating

Chlorine-36 is produced naturally in the stratosphere when atoms are bombarded by cosmic rays, high-energy particles that streak through space from beyond our solar system. This so-called cosmogenic production of chlorine-36 has varied over time due to fluctuations in the strength of Earth's magnetic field. When the field is weak, more cosmic rays can reach the upper atmosphere and more chlorine-36 is produced. Moreover, because Earth's magnetic field deflects cosmic rays toward the North and South Poles, chlorine-36 production rates decrease with distance away from the poles.

Researchers can determine the rate at which chlorine-36 was deposited in the past by determining the ratio of chlorine-36 to regular chlorine atoms in materials that contain chlorine and then comparing that ratio to the age of the material.

Beryllium-10 (10Be) Dating

Beryllium-10 is rare and forms from oxygen-16 and iodine-129 by cosmic bombardment of xenon. These isotopes can tell the ages of the length of time rock has been exposed on the Earth’s surface.

Aluminium-26 (26Al) Dating

Aluminium-26 decays to magnesium-26 (26Mg ) with a half life of 7.2 × 105 years. If this decay product is found in a solid, the 26Al/26Mg ratio will determine its age.



Relative Size Of Canadian Meteorite Craters


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