Isotope Evidence for Climate Change

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RONALD SMITH: Well, today I’m going to finish up a two-lecture description of paleoclimatology and the ice ages, and then starting on Friday, I’ll go into the subject of global warming. And I’ll be lecturing about that Friday, next Monday, and then you have your exam. I may even be continuing that subject on the following Friday. So we’re going to spend about a week on global warming. And it’s a very controversial subject, very interesting scientifically, politically, socially. And we’ve laid a good foundation for studying it. All the things we’ve studied up until this moment, including this stuff, I think will help us as we try to struggle with this issue of global warming. So any questions there? If not, so you remember, we started this last time. And I’m just going to flick through so these are the subjects we’re doing.

And I did these first three last time. Today we’re going to finish up the remaining four subjects within this general area. So just to flash everything by you. We talked about the last glacial maximum. We got here. So let me put a couple things up on the board just to reinforce what’s here. Hydrogen has a nucleus with one proton, therefore it has a single positive charge and typically has one electron with a negative charge going around it. There are two famous isotopes of hydrogen. One is deuterium, and one is tritium. I’m not going to talk about tritium today, but I’m going to talk about deuterium. It has one proton and one neutron. This has a mass of one. This has a mass of two. But notice it still has just one positive charge, so it has just one electron going around it.

So chemically, the two are identical. They have the same electronic orbitals, and they’ll bond and so on in just the way–in the same way that the other one does. The other one I want to think about is oxygen. Oxygen typically has eight protons and eight neutrons. Because of eight protons, it has eight positive charges. It’ll have eight electrons going around it. And that will determine its chemical properties. And this would be called oxygen 16. Sorry. Oxygen 16, because that mass adds up to 16 atomic mass units. The other one we’re interested in has 8 protons and 10 neutrons, adding up to a mass of 18. Still has eight protons, therefore has eight electrons. Same chemical properties, but a different mass, and this one is called oxygen 18. And by convention, that mass number is put to the upper left of the letter symbol for that element. So oxygen 16 and oxygen 18.

Now, we’re interested in water. So the question is what water molecules can we make from these isotopes? And by far the most common would be normal H2O, which would have oxygen 16 plus two hydrogens coming off it. Water vapor has a dihedral shape. It’s not a linear molecule. The hydrogens droop off it at an angle. That’s by far the most common, because this is much more common than that by a factor of 1,000 or so. And this is much more common than that by a factor of 1,000 or so. So if you just randomly put these molecules–these atoms together to form molecules, almost everything is going to be like this, but there are going to be some that have this structure where you replace one of the hydrogens with a deuterium. Notice that this has a mass of 18, because it’s 16 plus 1 plus 1. This is going to have a mass of 19, because it’s 16 plus 2 plus 1. And then the other one is where you replace the oxygen 16 with oxygen 18 but keep both of these as hydrogens. That’s going to be 18, 19, 20.

So that’s going to have a mass of 20. Now you can imagine others. For example, you could replace both of the hydrogens with deuterium, but that would be so rare that I’m not interested in it. Or you could replace the oxygen 16 by 18 and one of the hydrogens by deuterium, but again, that is so rare that I’m not interested in it. So I’m going to just limit myself to these three isotopes. This would be called the normal, and these would be the heavier ones. And we keep track of the ratio. For example, if you’re comparing these in a sample, you would measure the D/H ratio. Or if you’re comparing these two, you would take the ratio of oxygen 18 to oxygen 16. And the higher this is, you would say–the common terminology is that the water is heavy if it has a large deuterium to hydrogen ratio, or it’s heavy if it has a large oxygen 18 to oxygen 16 ratio, or light if it’s got a lower than usual ratio of deuterium. So I’m going to be using these terms heavy and light in a relative fashion to describe whether these ratios are typical for ocean water or larger or smaller than that.

So I hope you’ve Googled or gone back to your high school or college chemistry book and reviewed this. Now, isotopes also appear some isotopes are so-called radiogenic isotopes. They are temporary isotopes that can be formed by nuclear reaction and then decay automatically with some half life. Those are called unstable, or radiogenic, isotopes. And maybe those are the isotopes you’ve heard about before, but I’m not talking about those. These are called stable isotopes, because they remain. They don’t have any natural decay. They just exist in these different mass forms, and they persist over millions and billions of years. So these are the stable isotopes–in this case, I would call these the stable isotopes of water, which is what I’ve labeled this slide. Any questions on that? So now for our purposes, it’s very important to know that the heavier isotopes, these two, have a slightly lower vapor pressure. Because they’re heavier, they move a little more slowly in the liquid, which means they evaporate more slowly than this, but they would condense more readily than normal water.

Therefore, when you do get a phase change, when you get vapor going to liquid or you get liquid going to vapor, there will be a fractionation. This is the other new word. A fractionation means a process where you change the ratios of these two things. For example, if you had a bucket of nuts in your room peanuts and cashews and your roommate happened to be a cashew lover, so when she was grabbing a fistful of nuts, she would tend to get one that had more cashews in it, that’s a fractionation process. What’s being left behind then is going to be a little bit heavier, a little more enriched in the peanuts and a little bit more depleted in the cashews, because she’s grabbing out her favorite nuts. That’s what’s going on when you’re evaporating or condensing water. You’re changing the ratio of peanuts to cashews by the way you’re doing that thermodynamic transformation, vapor to liquid or liquid back to vapor.

That’s what we’re going to be tracking today. Now, the way we keep track of these ratios is with the so-called called delta notation. We have to define a reference ratio of D/H. And that’s standard mean ocean water. It’s basically a sample of several places around the ocean. They collect, mix them all together, and store large volumes of it in laboratories, and use that as a reference every time they’re doing an isotope ratio measurement. So that’s called–the abbreviation for that is SMOW, standard mean ocean water. Then you measure the D/H ratio in the sample of interest, divide it by the D/H ratio in your reference liquid, SMOW, subtract 1, multiply by 1,000, and that is so-called the delta D value. If it’s deuterium, it’s delta D. If you do the same thing with oxygen, it’s the delta O-18 value. And that’s what we’ll be tracking.

So notice the way this is subtracted. So if you had a sample that had a larger ratio of D/H than the reference, that would end up being a positive number. And we would call that– we’d say the sample is heavier than SMOW. If the D/H ratio is less than the standard, then the number is going to turn out to be negative, and we’re going to say that sample is lighter than SMOW. This quantity is usually stated as per mill or written not with a percent sign but with a percent sign with two zeros on the bottom. So again, a reservoir of liquid water under evaporation will become isotopically heavier because the lighter isotopes will evaporate more readily. A reservoir of water vapor imagine some water vapor in this room, and we’re condensing out some of it that’s going to be isotopically heavier–sorry, lighter, because the heavier isotopes are condensing out and leaving preferentially. You can do these experiments in the laboratory. You could take a dish of water. Take a sample from it.

Put it into a mass spectrometer to determine the isotope ratios. Then let it sit there for a few days and partially evaporate. And when you came back and took another sample of the liquid that would be this case and again put it on the mass spectrometer to get the isotope ratio, it would be heavier–isotopically heavier, because the lighter isotopes would have evaporated preferentially. So it’s an easy thing to do in the laboratory to measure this physical process called isotope fractionation. Now, in the real world, if you start with an ocean value of zero, the delta D value, or the delta O-18 value would be zero, which is likely, right, because ocean is our modern reference for isotope ratio. So typically, the delta values for modern ocean are going to be very close to zero. When you evaporate, the isotopes–the lighter isotopes evaporate preferentially, and so the vapor you have in the atmosphere is typically going to be lighter than the ocean. It will have a negative delta value.

And then as you condense water out, the rain is going to be heavier than the vapor. So if the vapor is minus 10, the rain is going to be heavier than that. Let’s say minus 3. The remaining vapor after you’ve removed some of the liquid is going to be lighter still, because you’ve removed some of the heavier isotopes. Then if you rain again from that lighter vapor, the precipitation is going to be heavier than that vapor but lighter than the rain that fell out before. And the vapor remaining is going to be even lighter, because you’ve removed heavier isotopes of the vapor that you started with. So this is the progressive fractionation process. This is a little bit tricky. So if there’s questions here, let me go back over this.

Yeah. STUDENT: [INAUDIBLE] PROFESSOR: I’m sorry? STUDENT: What does the negative hydrogen mean? PROFESSOR: So that’s the delta value. That would be–iff that’s delta O-18 and I think those values would be appropriate that would be the value here. And if it’s negative, it means that that number was less than that number. So that ratio was less than one so that when I subtracted one from it, I got a negative number. So what it’s telling me is that that sample is lighter than my ocean reference, my SMOW reference. So all the numbers in this cartoon are delta values, and they’re indicated that by that little per mill sign there, that per thousand sign there. That means not per hundred like a percent, but per thousand. Question. STUDENT: So it’s compared to SMOW water sample that came from the ocean? PROFESSOR: That’s right.

That’s why this first value is zero, because it’s being compared with ocean. And so that delta value turns out to be zero if it is ocean water you are sampling. STUDENT: So what’s the difference between the rain and the precipitation? Why is one lighter than the other? PROFESSOR: Well, the rain–this is the rain. I hope I didn’t misspeak. Rain and precipitation I’m using synonymously. But I’m drawing a stark contrast between the vapor and the rain. So here the vapor you’re starting with is minus 10 per mill. The rain that comes out of that, while it’s lighter than ocean, it’s heavier than the vapor from which it’s formed. So it has a negative number, but not as negative as that, making the vapor still lighter, making the next precipitation still lighter. And the vapor is even lighter than that vapor. So the things progress along together, gradually raining out the heavier isotopes.

Yes. STUDENT: Why are evapotranspiration and the vapor–the second vapor–different numbers? PROFESSOR: I’d rather not talk about this one, because it involves a complicated process of what water has fallen to Earth and what is being re-evaporated. So let me just ignore that. This is a correct value on average, but I’d rather not get into that, because it complicates the picture. We’re just going to deal with primarily these issues here. Yes. STUDENT: Does the percentage of the vapor as it gets increasingly smaller, or more negative, how does that affect the weather? PROFESSOR: Not at all. This is a climate diagnostic subject. It has nothing to do it all with how the atmosphere works. That’s the thing about these isotopes. They’re behaving chemically like ordinary elements, and this slight shift in the vapor pressure is so small that it doesn’t have any known influence on the weather patterns. So we’re just using this as a tool to understand what nature is doing, not as something that will change nature.

So when you take a–when you go to Antarctica or Greenland and dig down and pull out an ice core, you’re going to get–you’re going to look at things as a function of depth. I’ll show you some pictures in just a minute. But from that ice core, you’re going to shave off ice from different depths and determine the delta D and the delta O-18 values as a function of depth. Generally, the isotopes in the ice cores are lighter than the ocean. We saw that, because for example, the precipitation falling on this mountaintop is lighter than the ocean. It goes back to this initial evaporation. The vapor became so much lighter as it evaporated that that lightness still remains, is even amplified by precipitation that’s occurred between the ocean and the location where you’re getting your sample. So the isotopes in the ice core will be lighter than the ocean, but they’re not going to be constant. They’re going to vary with climate.

And in the literature, there are two ideas that have been put forward to understand why those isotopes would vary with climate. One is that under a different climate situation, the source of the water vapor might change. Something back here might change. And the other is that the fractionation that takes place in the atmosphere between the ocean and your sampling location might change. So I put that here. Source changes or changes in that progressive fractionation in the atmosphere. It’s believed now by almost every author that this is the dominant one. So I’m going to focus on this. So in the next bullet, I said it’s mostly sensitive to local air temperature, as that measures the degree of fractionation, which means that you get a light isotope signature during ice age.

I’ve tried to illustrate that here. So you have some water vapor source. We’re going to presume for the moment this is constant, giving you the same isotope ratios in the evaporating vapor. Then as this air parcel moves along and encounters different clouds and different storm events, you’re going to gradually precipitate out that water. Every time you do so, you’re going to be precipitating the heavier vapor, leaving the lighter behind. So the next precipitation is going to be lighter than the first and still lighter and still lighter. And so the isotope ratio in this final falling snow that’s going to fall on your ice sheet is going to depend on how much you’ve cooled the air between the source and the sample area, because the more you’ve cooled it, the more water vapor you have removed. Remember, when you drop the temperature, you decrease the amount of vapor that can be held. And therefore, if you cool the air a lot, you’ve had to remove more and more of the water vapor.

And when you do that, you get an isotope which gets lighter and lighter in proportion to how much you’ve cooled the air going from source to record location. All right. So we’re going to look at some Greenland cores. There have been a number of them done. Camp Century, a couple of GRIPs, GISPs, GISP 2, and Dye 3. Here’s a snowfall map of Greenland. Most of the accumulation is down here in the southeast, but you get some accumulation everywhere. And what we’re going to do is to put a drilling rig up on the ice cap, up on the ice sheet and drill down. Here’s a scientist marking an ice core. And ice core is about that big around. And it can be many hundreds of meters, even a kilometer or so deep. And of course, they pull it out in sections. They pull out a section break it, lay it down, put it in a storage shed, then drill out another ice core and so on. So they end up with these things stored in sections, and then they can do analysis as a function of depth.

And if they have some kind of a dating tool, they can determine what age that ice is at, of course older and older as you go deeper down into the ice core. We’ve done this on Antarctica as well. All these dots are places where there have been ice cores. We’re going to be talking about the Vostok Core, which is one of the deepest cores. This one was done by the Russians some years ago. So here’s a record then with years before present on the x-axis. And in the upper panel are cores from Antarctica. And they are represented as delta D values. Here they’ve written it as delta 2H. Remember 2H is what we’re calling deuterium. So this is the same as our delta D value. And then down in the lower panel is the delta O-18 for a Greenland ice core in fact two of them, NGRIP and GRIP.

And you can use either isotope. It doesn’t really matter which pair you’re using. You get a similar signal for climate change. There isn’t a great deal of difference between using delta D and delta O-18 for these purposes. But the numbers are different. For example, for delta D, the delta values are much larger. They’re minus 460, 440, 420, whereas for oxygen, which fractionates less strongly, you get somewhat smaller. Notice that they’re all negative though, which means that all these ice samples are isotopically lighter than the ocean. We expected that for two reasons. One is the water that evaporated from the ocean was lighter than the ocean water itself because of fractionation. And then this path fractionation as the water moved from source to Antarctica or source to Greenland, you’ve made it even lighter by raining out a fraction of the water on the way to this final record location. Well, this gives us a rather nice time record.

This is the modern era. This is today, if you like and it looks like we have somewhat heavier isotopes than normal during the last 12,000, 15,000, 20,000 years ago. And of course, this is the Holocene Period, the last interglacial. Prior to that, the isotope ratios were more negative, which means lighter. And we would interpret that–according to the dominant literature on this–we would interpret that as having more water removed between source and sample location, which means a lower temperature at the sample location. So the conventional wisdom is to interpret this as giving you a temperature record near the place where the record is kept, near the ice sheet. So colder. Staying cold for about 100,000 years. Then a brief interglacial, and then back into another long ice age with cold conditions at least at the location of the ice sheet.

May be cold other places as well, but what we’re measuring here, we think, is the temperature of the air over the ice sheet where those snowflakes are being formed. This is Antarctica. And notice that both cores done independently show a very similar record. Even more surprising is that when you go to the northern hemisphere in Greenland, you get a record that, again, is fairly similar with interglacials, glacial periods, the Holocene. So this seems, although it’s a local measurement, it seems to be global in character. And this is a characteristic then of global climate, even though it’s a local temperature record. Questions on this? I still find this astounding that these different ice cores from different hemispheres give such similar records. Yeah, Jordan STUDENT: Can you explain again why the lighter isotopes means colder temperatures? PROFESSOR: Yeah, because in this model, let’s say this is at a fixed temperature here. And so if you just cool the air a little bit between source and ice sheet, you would remove maybe half the water vapor and fractionate it accordingly. But if this were much colder, you would remove a much larger fraction, maybe three quarters of the water vapor, that would fractionate more strongly, and you’d get lighter isotopes falling on the ice sheet.

So temperature, remember, controls how much water vapor you have in the atmosphere. So the more cooling you get, the more water vapor has to be removed. The more water vapor you remove, the more fractionation there’s been in the isotope ratio. OK. So it’s remarkable that these two local records of fractionation and temperature are so similar. They provide an accurate time record of the ice ages. One more thing that’s even more astounding, you can take those same ice cores and find tiny bubbles in them. As the snow fell and more snow fell on top, and you finally compacted it into ice, there are little bubbles that trap samples of the ancient atmosphere. You can stick a probe into those ice bubbles and do a carbon dioxide concentration analysis to find out what the carbon dioxide concentration was in those ancient atmospheres.

So here I’ve plotted up from the Vostok Core alone, a record that goes the time axis is switched here. This is today, and this is going back in time. This goes back a bit further. I’ve shown you data back to 400,000 years ago. Here’s the isotope record converted to a temperature. Isotope record converted to an equivalent temperature. And then in the top panel is the carbon dioxide concentration taken from the same location along the ice core. Here’s the modern day. At the top of the ice core, you had about 280 parts per million by volume, which is about the value we think is the pre-industrial carbon dioxide value for the Earth’s atmosphere. Then back in the end of the last ice age, the last glacial maximum is here. You had much lower carbon dioxide concentrations, even below 200 parts per thousand. It stayed low, and others have mirrored the temperatures quite remarkably.

Carbon dioxide concentration and temperature are very similar. They have a very similar record, at least over this period we call the Pleistocene, with the ice ages coming and going and coming and going. We’ll talk a lot about this coming up in the global warming, but at a very minimum, this reinforces our idea of the importance of CO2 and climate. When you had high CO2, you had warmer climates. When you had low CO2, you had colder climates, perfectly fitting with our idea of the greenhouse effect and CO2 being a greenhouse gas. Now there’s also some argument though about cause and effect, which comes first, which leads and lags. We’ll get into that a little bit next week. But for the time being, I would take this at face value and just say that it reinforces the connection between CO2 in the atmosphere and global temperature.

Now I want to talk about other kinds of sediments. On land, you can walk around and see rocks like this. I’m sure you have. Rocks that seem to have a layered structure. Most of the time, when you see such rock patterns, you’re looking at some kind of evidence of ancient climate change. One kind of climate allowed sediments in a lake to look like this. A change in climate changed it to that. A change back gave sediments like this. This is easy to speculate about, because you see lots of sedimentary layered rocks like that. It’s not so easy to figure out what exactly this means in terms of climate. But it’s a start to be able to see such stratification on land and just imagine in a vague way, this may be evidence of changing climates in the ancient world. How many have been to the Grand Canyon? You see something like this. And lots of layered rocks. And what you should be thinking of when you see rocks like this is ancient climate change. These changes probably are telling us something about changes in the ancient climate of the Earth. But again, they’re rather difficult to sort out, because we don’t always know the environment in which those sediments were laid down.

So we’re going to do something similar to this but in the ocean, where the ocean sediments have had a rather constant sedimentation rate. And therefore, we can do a little easier job of interpreting if we can get a sample of that. So the way we do that is to do deep ocean drilling. You take a ship that has a drilling rig on it. And you lower that drill bit down to the ocean floor and then start pushing into the soft sediments. And then after a while as you go deeper and deeper, you’re drilling into rather solid rock, but it’s all representing deep ocean sediments. And they look like this. They’re about the same size. But they’re made of rock. And here’s an archive of hundreds and thousands of such deep sea cores that have been drilled out and stored. So we have a lot of these samples. We can do a lot of work with them. And primarily these deep sea sediments are comprised of calcium carbonate. They’ve got other things in them, but calcium carbonate, CaCO3, it has oxygen in it. Besides calcium and carbon, it has oxygen in it. And they’re mostly the remains of shells produced by microorganisms, phytoplankton mostly, in the ocean.

And when the organism makes its shell, it gets its oxygen from the oxygen in water. So there stored in this calcium carbonate is a record of the isotope ratio of oxygen– O-18 to O-16 in ocean water, because the organism got its oxygen from seawater in order to make that shell. I’m going to show you a deep sea sediment core, RC13-259, the one right there. Here is Antarctica, Cape Horn, Cape of Good Hope, Australia, New Zealand. And the deep sea sediment core in the ocean is here. And we’re going to compare that with the Vostok Ice Core, which is there. So what we’ve done is take samples of the oxygen from the calcium carbonate. We expect that sample of oxygen is going to be slightly heavier isotopically than the ocean water, because when the organism makes its shell, it slightly prefers the heavier isotope.

In addition to that, though, we expect it to change with climate either because the temperature effect on how the organism makes its shell might be coming through, or the isotope ratio in the water itself might be dominating the signal. The first papers on this thought it was this. Now there’s an almost universal agreement in the scientific literature that when you get oxygen 18 data from ice cores, you’re mostly getting a measure of the oxygen in the ocean water, and that’s telling you something about the glacial ice mass, how much water has been stored up on land in the glaciers. So most of the literature now is purporting this idea, that it’s most sensitive to continental ice mass. That means it’s going to be heavy during an ice age. Let me explain that with a cartoon. In a period like today, an interglacial period, the oceans are full of water.

This is our reference state today by the way for O-18. During a glacial period, you have lowered sea level by evaporating water from the ocean and not returning it. Remember the Quinnipiac field trip, right? You’re evaporating water from the ocean. You’re raining it out over land, and it’s running right back into the oceans in rivers. No net change. The hydrologic cycle is just running in steady state. The oceans are staying at the same level and so on. However, during a glacial period, you’ve evaporated some of that water and not returned it. You stored it as a large sheet on the continents. We know that that water is lighter isotopically than the ocean, so you stored light water on the continents, which means that the oceans are now isotopically heavier than they were before. And that’s the signal we hope to see in the deep sea sediment cores. So here’s the comparison. Vostok delta D is plotted in the upper panel, and the scale is on the left.

This deep sea sediment core, which is represented as delta O-18, is plotted on the right. And it’s the lower panel. And the time scales are matched up. So we can go back 400,000 years into the past for both of these records. And notice how they line up. But the signals are opposite. So during an ice age, the delta D is more negative, whereas the deep sea sediment core is more positive, because you’ve stored light water on land, made the oceans heavier, and that signal is coming through. So you see, this is perhaps even more remarkable, because we have two very different kinds of records, a deep sea sediment core and an ice core, whose isotopic signals are driven by somewhat different physics, and yet they both give us a corresponding similar record of the advance and retreat of these ice ages.

It’s a rather remarkable result. That’s why these isotopes are so valuable in giving us a quantitative understanding of the Pleistocene Period. Any questions on that? So you’ll have to think about this, because I know it’s a complicated subject, but these notes are posted. So part of the story then and this is not a surprise if you put all that ice up on land, you will have lowered sea level. And so here is a record from physical measurements, going around various places, understanding from the nature of the shoreline where sea level was at various times in the past. This is thousands of years ago. And at the last glacial maximum, which is here, sea level was about 120 meters below where it is today. And you find that same thing no matter where you measure. Remember, sea level tends to be level. It’s in the very name. And so when you drop sea level, you drop it the same everywhere in the ocean. And therefore you should get the same record no matter where you do that analysis. And then as we came out of the last ice age, we came up to the present sea level rise.

So this is from physical measurement, not from isotope measurements, but it agrees conceptually with that deep sea sediment core analysis, which I argued is a measure of how much water is stored on the land in large ice sheets, especially the Laurentide ice sheet and the Fennoscandian ice sheet. So if you made a map of the Earth during the last glacial maximum, land would be somewhat larger. Of course we know this is mostly continental shelf. You don’t have to lower the ocean very much in order to expand it out a bit–expand the continents out, because there’s some shallow submerged land today. But that’s typically where the coastline would have been 15,000 to 20,000 years ago. And the point I want to make here is that the gap between Asia and North America was actually bridged at that time because of a lowered sea level. It’s believed then that this was the moment when humans could first migrate to North America. And so this business of sea level change and ice ages then directly carries over to the anthropology of the Americas. When did people first come to this continent, and how did they get there? And there are other theories.

They may have sailed from Asia. They may have sailed from the islands in the Pacific. But this, I believe, is still the leading theory for how humans, which evolved in Africa and Asia, finally made it to the New World. And they did that about 20,000 years ago following the Bering Sea Land Bridge. The way anthropologists, or archaeologists describe the different cultures in the Americas is by the type of arrowheads that they made. So the Clovis culture sorry about that would have made the Clovis points. They can track this back to about 11,500 years, but before that, there’s not too much record of humans, which means probably that’s about when they came. So roughly during the period of the last glacial maximum is probably when these cultures moved to the New World. And all the Indian cultures we see today then in the desert southwest or up in the northwest US and Canada, those cultures probably came over about that time. There is a national park up on the northwest coast of Alaska.

I want to go there someday. It’s called the Land Bridge Preserve. And on their website, they show human footprints, which supposedly are a record of those people walking It must have been cold though, but I guess they had built up a resistance to the cold over from Asia to North America during that period of time. I only have a couple minutes to talk about Milankovitch theory. Your book does a pretty good job of it, but I do want to spend a couple minutes on it. What has caused this periodic coming and going of the ice sheets during the Pleistocene? It’s not a solved problem. I don’t want to present this to you as known fact, but there is considerable evidence that slow changes in the Earth’s orbit induced by Jupiter and other planets probably were what caused or at least in some way paced the coming and going of the ice ages during the Pleistocene. And the three changes in the orbital parameters that are discussed–that are understood in this regard are the precession that is a slow rotation of the orientation of the Earth’s tilt.

Second is the obliquity, which is the magnitude of the Earth’s tilt. And the third is the eccentricity of the Earth’s orbit. You know that the Earth’s orbit around the sun is not exactly circular. It has a little bit of eccentricity to it. And the amount of that eccentricity has been changing slowly with time. The precession goes through a full cycle all the way around and back in about 20,000 years. The obliquity, this little changing of the tilt, goes through a full cycle approximately every 40,000 years. And the eccentricity, the changing of the ellipticity of this orbit, goes through a cycle of about every 100,000 years. This is the smallest effect. And recent papers suggest that this may not be playing much of a role. Most of the attention the last few years has been on these two cycles for causing these advances and retreats.

What these things mostly do, if you think about it, is that they change the way seasonality works. Remember, the obliquity, the tilt of the Earth’s axis, controls seasonality on Earth. And the precession, the way that this orientation of the tilt lines up with the long axis of the orbit, controls the strength of the seasons also because of the way the tilt and the perihelion combine. So these two things are mostly changing the seasonality, but apparently that is enough to push us into and out of an ice age. This diagram shows the same thing in different order, the eccentricity, the obliquity, and the precession. And I think that’s all I have to talk about today. Yeah. So we got just through this at the end. I think you have enough to finish up that problem set. And on Friday we’ll begin to work on the subject of global warming. Read chapter 13. Chapter 13 is your key in the book.

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