9UNP1290HD_Clark_Johnson_WordPress

- Welcome everyone to Wednesday Nite @ The Lab. I'm Tom Zinnen. I work here at the UW-Madison Biotechnology Center. I also work for UW-Extension Cooperative Extension and on behalf of those folks and our other co-organizers, Wisconsin Public Television, the Wisconsin Alumni Association and the UW-Madison Science Alliance, thanks again for coming to Wednesday Nite @ The Lab. We do this every Wednesday night, 50 times a year. Tonight it's my pleasure to introduce to you Clark Johnson. He's a professor in the Department of Geoscience here at UW-Madison. He was born in San Francisco, raised in Palo Alto, went to UC-Davis for his undergraduate and got his Ph.D. in Stanford, and then came here to major in ice fishing. 
(audience laughing)

 He's been here for 30 years. It's a pleasure to have him back here at Wednesday Nite @ The Lab. Recently, NASA confirmed the discovery of liquid water on Mars.

Clark has been instrumental in astrobiology for a long time. Today we're going to hear about water on Earth and how bacteria pulled iron out of that water and laid down iron ore, but it's going to be pretty interesting to hear what we learn here on Earth and how it might play out on Mars. Look forward to hearing more about it from Clark. Please join me in welcoming Clark Johnson to Wednesday Nite @ The Lab. 
(audience applauding)

- Well, thank you very much. Tonight we're going to go back in time. 

I know to a lot of people a million years sounds like a long time. We're going to be working on the billion year time scale. One of the things that I really like about working on early Earth problems is that none of us were there, as far as we know, which means that we need to develop very clever ways of teasing out what the surface of the Earth was like, what life was there. And as Tom mentioned, it's really important to study things like the early Earth, because that informs us about other planets and vice versa, like Mars certainly is in the highlights quite a bit. I'm going to talk about really two papers that came out from our group this year. And I think the one that caught Tom's attention was this one that came out in the Proceedings of National Academy of Sciences back in July with one of my post-Docs, Weiqiang Li. And there was a nice US or UW-Madison news report on that. And then just yesterday actually, we had another news release. And this paper came out on Earth and Planetary Sciences, which we think nails down the origin of oxygenic photosynthesis. So I'll be kind of blending these things in, and the common theme of that will be: what is the role of iron? 

Tom mentioned that I'm an astrobiologist. I'm trained as a geologist, but over the years I've worked more and more with NASA. The work that I'm going to talk about was funded by the NASA Astrobiology Institute, the webpage there on the top one. And then we have an Astrobiology program that we started in 2007. I'm the PI of that, and that's covered a wide range of research topics, just one of which I'm going to talk about. Since 2007 that we've started, we've had 35 students and post-Docs graduate from this program, if you will. Some have gone on to become faculty members. Some are actually leading NASA missions, like the Mars 2020 mission. And others work for oil companies. And you might kind of wonder why does an astrobiologist end up at an oil company, and I think it's a nice testament to, you know, thinking out the box, thinking outside the box. 

That's what we want for our graduates. You know, bring in multiple lines of disciplines. Think about programs like the early Earth, and I think that's attractive on multiple levels. So this is briefly what I'm going to cover. I want to first talk about the importance of iron in the universe. You may not realize that actually it's kind of an unusual element on the Periodic Table in terms of its abundance. The importance of iron in biology, and this is even more important when we look at early biology, relative to, let's say, biology today. People often ask me, "How do you know which rocks to go find?". And so I want to spend a little bit of time about how we look for evidence for life in the rock record. Then we're going to get into a little bit of geochemistry. 

I'm going to keep it very simple, but we're going to talk about things like isotopes. We'll have to talk about a little bit about electrons, but we'll keep it simple. And then I'll end talking about iron formations, something that has been very important in the Superior, Lake Superior region economy and turns out to be really important to understanding early biology. So let's start first. What is the importance of iron in the universe? This is a chart that I know may not show up too well, but it's a plot of the abundance of the elements on the y-axis versus the atomic number. And in the upper left corner there is hydrogen, which is the most abundant element in the solar system. Next is helium, reflecting the fusion of hydrogen to helium in our star. And then you see this kind of logarithmic drop. But if you kind of look in the middle there and you see iron, iron is unusually abundant for its atomic number. 

It defies, it's kind of this spike, it defies that downward trend. What is the reason for that? Iron is, we've got a lot more of that than the general abundance would show. And the reason for that is a mass problem. And some of you probably know this. I'm going to bring something that I know a lot of people dealt with when you were in introductory chemistry classes and you had to calculate the weights of all your reagents and things. And you looked up that weight of oxygen, right, and it's 15.999, a little under 16. If we add up the sum of the individual parts, however, we would come up with the mass of a proton, which is greater than one; a mass of a neutron greater than one; and we actually end up for the major isotope of oxygen has a mass of 16.12, if we added up the individual components. So what's going on? Where is the missing mass when actually it's 15.99? 

It gets even worse when we consider multiple isotopes of oxygen. So you remember the number of protons defines the element, but we can add different amounts of neutrons, and there are actually three stable isotopes of oxygen. Sixteen is the one on the left, that's the most abundant. Then there's 17 and 0, and then 0 18. And so this mass of 16.12 gets even greater than the 15.99. What's going on? This is a plot of something called the binding energy. And so Einstein told us that energy and mass are interconverted, and we can have his expression, you all remember that one from school, E=MC2 or Delta M, Delta E equals Delta mc squared. And this curve that's going up towards iron is constantly increasing. So from hydrogen to iron, Delta E is positive and so the energy we get by converting mass into energy is giving us increasing energy as we go on. 

In other words, nuclear fusion in a star is exothermic; it releases more energy the higher up we go. But that inflection up there occurs at iron. Then notice that to the right the curve starts tapering down. We lose energy. And so basically, iron is the end of the line for nuclear fusion. All paths must lead to iron. And it's kind of like if you started rolling down these marbles, down one side. They would all accumulate at the end, right? And that's why iron is unusually abundant. Now we are not making iron in our star. 

We're only making hydrogen to helium. We need truly massive stars to make that, and we need very high temperatures. If we kind of look at a plot here, this is a plot of luminosity versus temperature. And if we look at that main sequence as we go from the lower right, which are very old, small mass stars, up to the upper left which are very large, young stars, stars on the order of tens of millions of years, far younger than ours, it's only the big ones that make lots of iron, and of course, the elements that we need for life like carbon and oxygen, nitrogen. All of those are made not by our stars, but by other star systems. Now the only way that we access elements heavier than iron, and in fact how we spread elements out generally in the universe, is when we have a supernova reaction. Now if we look at the terrestrial planets, the inner rocky planet, all of these have iron cores. Most of the iron in the Earth is actually in the core. The iron and all the carbon, oxygen, nitrogen needed for life is actually on other stars, stars that are very large, very short-lived, probably too short to develop life. Tens of millions of years is probably too short, and that's kind of when the iron is. 

The star systems that develop the elements for life cannot support life because of their short lifetime. But we can inherit those elements into longer-lived solar systems, like our sun, which is more intermediate in mass. So iron is special in the universe. It's unusually abundant, and there's actually a very good physical reason for it. Now what about this question about the importance of iron in biology? So one of the things that we have to do is we have to go and think about how iron was different in the Earth's past, relative to today. And biology took advantage of that. We need to talk a little bit about the chemistry of iron. And when I mean chemistry, you might recall from chemistry classes that means moving electrons, right, making bonds. And so we talk about oxidation of iron, right? 

We all know that if you put a big pile of nails out in a bucket of water out in the atmosphere, it'll turn to rust. There's three oxidation states of iron. The picture on the left there is an iron-silicate meteorite and the iron parts are the silvery part. That's pure iron metal which has a charge of zero, hence that Fe0. The yellow colored ones on the left figure are the olivines. That has a charge of two. And then of course I just, you can look at some rusted chain. That has a charge of three, and so that would be ferric hydroxide, would be the chemical formula. Now today's Earth's interior is the one on the left, iron zero, iron plus two, what we might call reduced iron. The surface of Earth today is oxidized iron, right? 

We see that in our everyday life. And that process involves a loss of electrons bonding with something else, some other reactions. And that is the key. Iron is wonderful at moving electrons. And if you were to distill down one of the most fundamental biochemical processes that life does, is it pumps electrons. It pumps electrons to make gradients in chemical potential to then synthesize biomolecules and drive energy. So, we're going to keep that in mind. Now what about the oceans? Where's all the iron coming into the oceans? They come from these hydrothermal systems that occur at what we call mid-ocean ridges, where two plates are coming apart. 

And you've probably heard about plate tectonics, and we have a lot of water. We have a heat source to dissolve iron, and that comes out as exhalative deposits. And these mid-ocean ridges are everything in blue that is kind of seeing these ridges there. Those are underneath the oceans. The largest mountain ranges on Earth are actually underneath the oceans. There are these long mid-ocean ridges. And this is what one of these looks like. This is hydrothermal vent with some tube worms. And then that plume is aqueous Fe2. The question then is: how does that aqueous Fe2, if that was really abundant in the early Earth, how might that have been used by biology? 

Now let's talk a little bit about a few biological processes. Everybody's heard of photosynthesis. You may not realize that there's two types of photosynthesis actually. Oxygenic photosynthesis, where the basic driving force is the sun, is moving electrons in this pathway. And the little guy on the left there, which is water, is called the electron donor. And you can see that little red electron will hop in. It will take CO2 and reduce it to organic carbon. And the molecule on the far right, that CH2O, is a very simple form of organic carbon. The waste product of that is the oxidized donor, which is O2, and so we call it oxygenic photosynthesis because, like we have, might have some algae, we might have some trees, the product is releasing oxygen. We use that for our respiration. 

We use aerobic respiration which is a very high-energy yield. There's an awful lot of energy locked into that CH2O, that organic carbon, which is fundamentally fueled by sunlight. If we look at the tree of life, and this is a little bit different tree of life here having a common point that some of you may have seen. The bacteria are on the upper left. The right is the archaea. Both of those groups are prokaryotes, or single-celled microbes without a nucleus. And the eukaryotes, which is what we are, is on the bottom part. And shaded in blue there are clades of life that are either hyperthermophiles or thermophiles. So early life likely was associated with hydrothermal vents, good source for the iron that I just showed you that, and the most deeply-rooted bacteria arcyria are hyperthermophiles or thermophiles. Now there's an earlier form of photosynthesis that based on looking at this tree of life, we would place more deeply-rooted towards the last common ancestor. 

And that's called anoxygenic photosynthesis. It survives today in the purple bacteria. And look over here, the electron donor for oxygenic photosynthesis was water. In this case, it's ferrous iron: Fe+2. Ah, okay. Early life really was probably using a lot of Fe2. And the product, the oxidized donor is not oxygen, but now Fe3, or the oxidized form of iron. We just kind of rearranged that. I'll put oxygenic photosynthesis on the top so the product is oxidized Fe3 and organic carbon on the right. And there's another process called microbial iron reduction that actually takes that Fe3 and that organic carbon and runs it in reverse. 

Those are just one reversal of the other. So both of those yield energy to life. Here's a little picture of some iron reducers. They actually pump electrons to ferric iron Fe3 oxides with these little nanowires. That has actually been imaged. You can drag a little probe over that, and you can measure a resistance or a voltage change. They're clearly electrically conductive. If we look at the tree of life and we add up all the ones that either reduce iron or oxidized iron, we get these guys in red. And it's quite a bit. It's found in the bacteria and in the archaea. 

And one of the really big questions then: is what are the different ages of these branches, okay? This is based on RNA, the relative mature, or the evolution of RNA. But what we really want to know is, can we put ages on those points? And that's where geologists come in saying, "Ah! Can we find a way of looking "at these metabolisms in the rock record "and then finding where we can start "putting time points in that?" So the early Earth was very abundant in Fe2 prior to the rise of oxygen which happened with oxygenic photosynthesis after early and oxygenic photosynthesis. And if we start looking at some of the most fundamental properties of life, such as nucleic acid, we find that there's an important role for iron in the past. So you all heard of RNA and DNA, certainly DNA. RNA performs many functions. It can also store information, in fact it may have been a precursor to what we call the DNA world. 

It's called the RNA world. RNA is also very important in making various biochemical reactions by joining molecules or splitting them apart. And the major role for a plus 2 ion today in RNA is magnesium. And that's involved in folding and RNA function. But a really exciting discovery comes from one of my friends and colleagues, Loren Williams at Georgia Tech, shown here with one of his post-Docs. And he said, "Well what if the, "if the geologist such as myself is telling "there's a lot of ferrous iron, "reduced iron in the oceans in the past, "maybe actually RNA was, "was actually started in iron "and then later that had to change." So he did a whole bunch of computational models, and here is showing a reactive center and RNA for magnesium on the left and one for iron on the right. Iron is beautifully stabilized in RNA. Today is on the left, where magnesium plays that catalytic role. But could it have been iron 2 in the ancient world? 

This is a plot of reactivity of magnesium-based RNA on the top, iron-based RNA on the bottom, relative to the nucleotide position. And basically, they're mirror images. Iron 2 is just as good in RNA as magnesium. It actually gets even better. If you look at the reactivity of iron-based RNA to either join or ligate things on the left or cleave or break things apart, you can see that it's actually much more reactive in magnesium RNA. And Loren Williams like to say, "Yeah, iron-based RNA is RNA on steroids." So he believes that the first RNA was actually iron-based, and then it lost that ability. So why would, if there was an iron-based RNA, why would it evolve to magnesium RNA? That's what we have today. What would happen? 

Well, let's kind of think what's going on. Let's go back billions of years ago. And we have the sun and we've got this hydrothermal source of aqueous Fe2. Then we have oxygenic photosynthesis, eventually rise on the planet. And what does it do? It makes oxygen.

But if you react that oxygen with that Fe2, you're going to have this reaction which makes insoluble iron3 hydroxides. It rusts out, literally. So once the rise of oxygen was sufficient, the availability of soluble iron went away. And the biology has actually probably been irreparably changed. 

So how are we going to look for this? Where do we go? We don't run around and just pick up random rocks hoping that there'll be something in there. Like any science, geologists build on decades upon decades of previous work by other geologists who have mapped other terrains, who exchange data with each other at meetings or in publications. That is key. This is a map of different age terrains, and I want you to focus on the blue and the purple. Those are the oldest ones. Those are the regions of the Earth that have rocks over 2.5 billion years ago. If we're in our own part of the world, it's what we call the Superior Province just north of here, or actually is a chunk of this old crust in Wyoming. The oldest rock in the United States is this guy, the Morton Gneiss. 

Some of you may have heard of that. If you ever went up to St. Paul and you saw this statue of Leif Ericson out at the state capitol, that pedestal is built on the Morton Gneiss. It is the oldest rock in the United States, not the oldest in North America. There are older rocks in Canada. But it is a very high grade rock, very typical of old rocks here. It's been subjected to metamorphism at temperatures of 500 to 800 degree C. That is not my first choice for looking for ancient life. But that's what we have in North America. We need to go elsewhere that have not seen such high grade metamorphic effects, and one of them is the Pilbara. 

That's in Western Australia. And we can see things for the exact same age, such as this beautifully preserved stromatolite, which is a layered microbial community that was fossilized, not subjected to metamorphism. Ah! Now we've got something! Another place that we do a lot of work is what we call the Kaapvaal Craton in South Africa. And I'm going to show you some results from both cratons. So we go where we have to, where the oldest low-grade rocks are. Now one line of research is to look for what I call direct-evidence for life. And the most common one would be looking for microfossils. So, we talked about cyanobacteria and when we mentioned, well no, I didn't mention that. 

That's oxygenic photosynthesis, a typical one is cyanobacteria. There's a fossilized cyanobacteria beautifully well-preserved on the uppermost plot there. And that is only 750 million years old, hence it's so well-preserved. Here are the oldest proposed cyanobacteria, the 3.5 billion year old rocks, 3.5 billion years is 3.5 GA, giga-annum. I sometimes use those interchangeably. You can see those are much more poorly preserved. These are highly contested. A little bit younger, we might go and find some from South Africa. These are extremely large microfossils. The scale there are 50 microns, you're probably not calibrated, but that is much, much bigger than a cyanobacteria. 

These might even be eukaryotes. Nobody really believes it that age. Ugh, these are contested. Unfortunately, these are very few in terms of localities. And the oldest ones are highly debated. Another thing that people look at are looking at the remnants of a layered microbial community, and these are stromatolites. It's a little bit bigger; it's a little bit easier to find a large stromatolite than lots of small, fossilized bacteria. Probably the most famous modern stromatolites are in Shark Bay in Western Australia. It's a hypersaline bay; that's why it's, it keeps the animal life away. Here are some 3.4 billion year old stromatolites from Western Australia. 

And you can kind of see the domical shape in those. You can look at the conical ones on the left and then a domical one kind of there in the middle. Unfortunately, there are very few of these localities. If we look at all the localities that are older than 3 billion years, there's a handful of 'em. It's not a very big what we call footprint in the rock record. So another approach is what I call an indirect approach to look for ancient life. And the thing that we're going to focus on are the minerals that are maybe produced by things like this iron oxidation or reduction by microbes like anoxygenic photosynthesis, oxidizing it, or microbial iron reduction, reducing that. The footprint is bigger. We have much, many more rocks to look at. Minerals are more resistant to the degradation of metamorphism relative to, let's say, organic carbon that is easily volatilized away. 

And we can tell, oftentimes, specific metabolism. So that's what our group is really focused on. What are the minerals that we want to look at? Well, I've organized them from left to right, in terms of reduced versus oxidized. And you'll recognize some of these. So pyrite's there on the left; that's all iron2. Siderite, an iron carbonate is iron2. Magnetite is mixed: has a little bit of iron2 and iron3. And them hematite, or rust, is iron3. So there's a relative order of electron-rich, electron-poor on the right. 

And people have looked at rocks at the first order. Okay, which ones have a lot of hematite? Which ones don't? Which ones have a lot of pyrite, right? Pyrite would be reduced so, we wouldn't expect pyrite to exist on the surface of the Earth very long with an oxygen atmosphere. Hematite is perfectly stable. And people came up with this classic curve. This is the atmospheric curve for oxygen, relative to time. The timescale there is in millions of years. So 4,000 million years is four billion years on the far right. 

And it ends at about 1.5 billion. The scale is relative to today. And notice that it, there is a big rise there that has been called the great oxidation event at about 2.3 billion. And notice that the maximum though is about 10 to the minus two relative to today. This is not a lot of oxygen. If we were transported there instantly, we would die of asphyxiation. But it was a big jump relative to the anoxic world. And the evidence was looking at, let's say, those rounded pyrite grains. Those were deposited by a stream on land on the right. That would not survive in an oxygen atmosphere. 

And so the youngest rounded pyrite grains are about 2.5 billion. And then these are some oxidized lake deposits, that image on the left. That was deposited on land. That has a lot of hematite, and so that would be after the oxidation event. Now we've mentioned things like banded iron formations. Where do those fit? Iron formations are in that histogram with the oxygen curve superimposed upon it. And you can see over on the far right, very early in Earth, hardly any iron formations. And then a really big chunk at about 2.7 to 2.5 billion years ago, actually before the great oxidation event, right? And then it kind of tapers off as we move to the left, as we go younger and younger, and we're getting less of that. 

In our own necks of the woods, these iron formations have been really important for the steel industries and the automobile industries in the upper Midwest. And they range in age from about 1.9 to 2.7 billion years old, kind of in the middle of that big histogram where a lot of them are found worldwide. The biggest unit is here. This is in Australia, right? Everything is big in Australia. This is the Brockman Iron Formation, 2.5 billion years ago where it is mined. It's the Brockman Iron Mine. It's the largest iron mine in the world. They generate 200 million tons per year of iron ore. And it goes out in these massive trains going out to be shipped out of Port Hedland to the north. 

It is, I guarantee you, we all, all of us own iron from this iron mine, because this is the big one. If we looked back into the rock record for these iron deposits, this is the oldest non-metamorphosed one. It also happens to be in Australia. Again, Australia is this continent that escaped a lot of the metamorphism that North America did not. This is called the Marble Bar Chert. That's 3.5 billion years old, and you can see those iron-rich deposits. Now a lot of these places are out in the middle of nowhere. Here's the town of Marble Bar from a Google Earth image, and that's the only town on that slide.

Marble Bar is known as Australia's hottest town because it has the largest number of days over the year that are over 100 degrees, 154 days of the year. Yeah, it gets pretty scorching out there. 

These are some little bit younger iron deposits. These now are in South Africa, 3.2 billion years old, the Manzimnyama Chert. So this is what we have to deal with. And we're going to start looking at iron, I'm going to show you data later that look at iron formations stepping younger and younger in time. What can we tell about the oxidation, right? And also the microbial reduction. We want to see how those play off each other.

So to do that, we need to spend a little bit of time on some chemistry. And that's kind of unavoidable. And we're going to talk about isotopes. 

So I mentioned this briefly when I mentioned about oxygen, so many elements have more than one mass. They all have the same number of protons for a single element, but we can add different numbers of neutrons. And that changes the mass of these, and they respond in very subtle ways to bonding differences.

What we're after is to quantify how much oxygen was out there. It's not enough to say, "Okay, there's some oxidized iron "and some reduced iron. "Okay, how much oxygen, how do I quantify that, "how do I model that to actually get numbers "for the amount of oxygen that was out there?" Okay. So we're going to look at isotopes, different masses of different elements. And so for example, if let's say we have a whole bunch of aqueous Fe2 that comes from those hydrothermal vents in the bottom of the oceans, and I have a little bit of O2. And that oxidizes to this iron3 charge, hydroxide. 

But I could have a lot of oxidation, right? I could have a whole bunch of oxidation. I make a lot of iron hydroxide. And you might say, "Well why don't you go out "and just measure how much iron hydroxide "deposits are out there. "Add it all up, see what did you get." The problem is we can't do that because our preservation of the ancient rock record is so poor because of plate tectonics and metamorphism that I only have a few little localities to actually measure. So I need something that's not dependent on the mass of that iron, but tells me now much oxygen was involved in oxidizing what iron was there. So we need some measure of the amount of oxidation of aqueous Fe2. We're going to go to iron isotopes, and there are four isotopes of iron. The most abundant one there is iron 56 at 91.75 percent in abundance. 

The next abundant one is the one on the left, iron 54, and so we can talk about the isotopic composition of that iron by looking at the ratio of 56 to 54 iron. So that representation is this delta value on the y-axis over there, this delta 56. And if I start with aqueous Fe2 that's at zero, and I oxidize it to aqueous Fe3, that is increased, and it's 56, 54, or this delta 56 parameter. It becomes heavier, right, because it's enriched in the heavy isotope. Normal iron would have this lower ratio, the aqueous Fe2. The reason it does this is because of the bonding differences between iron 3 and whatever its bound to and iron 2 and whatever it's bound to. They have actually different vibrations. And those different vibrations substitute the heavy isotope into the high frequency species. The lighter isotope into the lower frequency species. Okay. 

So there's good theory behind this. If I take that aqueous Fe3 and precipitate it as ferric hydroxide, I get about the same result. But there's one really important thing that I need to remember: the extent of reaction is really important. So let's imagine that we start with aqueous Fe2. I put it in a closed box. I oxidize it 100 percent. What's changed in the isotope composition? Nothing. I'm back to where I started from. It's actually only the little bit of oxidation that shows me the biggest effect. 

And as I approach 100 percent oxidation, I'm back to where I started from. That means that under low oxygen conditions, I'll have the highest delta 56 values, or the highest enrichment of the heavy isotope. And at high oxygen contents where I have oxidation to near completion, I'll have low values. Okay? And we've calibrated this in the laboratory, as well as modeling.

There was a reduction process in the microbial story though, too, right? Remember that? We talked about microbial iron reduction, kind of running that in reverse. That actually makes very light iron. Now notice the scale on the 'Y' becomes negative. 

That means it's really, really light in the heavy isotopes, enriched in the light isotopes. And so that's the signal for microbial iron reduction. So I have a positive signal for a slight amount of oxidation. Essentially a zero signal for 100 percent oxidation. And then a light signal for microbial iron reduction. This is a couple of pictures. I thought you'd like to see where is done. It's just down the street in the Department of Geoscience. We do a lot of our work in a clean lab, because we work on really small samples. This is our oldest mass spectrometer. 

It's almost 30 years old. Still works well, not quite as nice as some fancy new ones. That mass spectrometer on the right, there, is about 15 years old. The blue box on the left is a laser oblation system that we sometimes use to sample very small samples. And then our newest instrument installed two years ago has a beautiful new detector system, allows us to analyze smaller and smaller things to higher and higher precision, right? That's what we always want, is shown right there. So what about these iron formations? That's the database that we're going to work for. And the two things we're going to try to pull out of this is the oxygen levels and what microbial metabolism was responsible. So let's go back to this Brockman Iron Formation, the biggest one on the planet.

Big problem, though: all of that iron is all surface weathering. The Australian continent has been exposed to weathering for at least 50 million years. And so the primary minerals that were laid down 2.5 billion years ago cannot be sampled in outcrop. So we do all of our work on new drill core that go below the weathering horizon, gets us fresh minerals. And when we do that, we recognize a big change in the iron formation mineralogy through time. I mentioned the Marble Bar Iron Formation, and in drill core and in outcrop actually, it's nothing but hematite. So all hematite, all Fe3. If I look at the Brockman Iron Formation, however, over here, I find a lot of Fe2 minerals. Yes, I have some hematite, but I have that siderite, that Fe2 carbonate, and I have some magnetite which also has Fe2 in it. Very different mineralogy when I look at it fresh in drill core. 

So now let's look at the data through time, and we're going to kind of remember what we think about in terms of how much oxidation and oxygen. And then when does that negative microbial iron reduction signal appear, and it's quite a stunning record. So the record starts at about 3.5 billion years old and moves to the left at 1.5 billion. And those high values in the upper right corner are what we calculate to be only partial oxidation. That's the Marble Bar Chert data. A little bit greater oxidation is that guy. Notice how that has dropped. And everything in red here only contains that hematite. There's no reduced minerals in there. The next set of data moving to the left includes some blue shales and also some green iron formations that also contain some Fe2, and so we think actually that big negative signal right there is where we've had the start of microbial iron reduction. 

Now in our PNAS, our Proceedings in National Academy of Science's paper that came out in July, we added the new thing in that was adding another isotope system that we don't have time to talk about that allows us to distinguish on that spread, that vertical spread, between a hydrothermalin member and a microbial member. We don't have time to work through those systematics of that other isotope system, but that's the bottom line conclusion of that paper. So how do we put this together in terms of what was going on on the Earth? Well you saw this diagram, but a little bit differently. I've separated them vertically so the classic atmospheric oxygen curve is up there on the top. And then here I've drawn kind of the ocean, the photic zone of the ocean, that little wavy stuff and all the blue below there, and that's the iron formation curve. And those are the oxygen contents that we calculate by modeling the isotopic data. So the oldest stuff, the oxygen in the photic zone, this is only in the oceans, was less than .0001 percent of that of today. That's essentially zero. But there's a big jump at 3.2 billion to about .1 percent of today. 

And that's where we put oxygenic photosynthesis, that big jump of several orders of magnitude. Notice as we go younger, we get about 10 to 15 percent of oxygen of today, but that's oxygen in the shallow water column where oxygenic photosynthesis is occurring. It's not yet oxygen that's made it up into the atmosphere. That big negative spike in the iron isotope composition that I showed you? It occurs right where the major iron formations are occurring. That we would put, the beginning of microbial iron reduction where that starts to go down. And then we would say that the peak was kind of that pile. That really negative iron isotope, that light isotopic composition of iron with the major iron formations on the planet, probably was when microbial reduction never had it better. It had all of the iron 3 oxide and the organic carbon from the photosynthetic processes. And that's the fuel that it needed to run the reduction, those reactions in reverse. 

Because we get most of our iron in our daily life from that peak right there, and that's a big microbial signal, I would argue that most of the iron we use was recycled by microbes 2.5 billion years ago, which is kind of an interesting thing. Those guys laid down the iron that we later drive around in. So the peak was about at 2.5 billion years ago, just before the rise of atmospheric oxygen. The oxygenic photosynthesizers are continuing to make oxygen. But until they fill up, all of what we call the reduce sinks, until all that iron is oxidized sufficiently, only after that can actually free oxygen accumulate in the atmosphere. And so that's why the major iron formation, that peak, actually just precedes the rise of oxygen. Now let's look at the data here. So the big peak is that negative one right there. It kind of, that negative signal kind of goes away, doesn't it? Why would that kind of decrease? 

And I kind of use the phrase, it decreased its footprint. On the other plot it looks like that. The iron formation deposition is declining. The oxygen is risen in this, in the atmosphere. And that iron isotope signal for iron reduction has dropped. There's another player that came onto the scene. When we oxidized the atmosphere, we had oxidative weathering on the surface of the Earth, it oxidized that pyrite. That produced an oxidized input into the oceans of something we call sulfate. And microbial sulfate reduction, which is a really important process in modern marine environments took over. The iron reducers got shoved off. 

They still exist on the planet, we find them, but in very small niches. The sulfate reducers are king because that's what's stable with oxygen. So what are the conclusions then?

Iron is unusually abundant in the universe for its mass. And it's important in redox reaction, it moves electrons. Because we go from iron zero, like this iron metal core that we might have in a planet, to iron 2 for the silicate Earth, and then iron 3 for the rust that we see. We can run those reactions in reverse, and microbes can take advantage of that. Because the fundamental process that microbes do with the environment is that they exchange electrons. That's why we talk about it in terms of electron donors and accepters. Prior to oxygenation of the oceans, however, iron 2, soluble iron 2 was much more abundant than it is today. 

The iron quantities in the oceans are vanishingly small. Why? Because all of that hydrothermal iron gets immediately oxidized and falls out around the vents. But before the oceans became oxygenated, there was a lot of iron in the oceans and biology actually probably used that as one of its core cation components. It was only forced later to abandon iron and use something like magnesium in the case of RNA shown here, because the iron 2 was gone. So a fundamental change in biology. Not, we have to think kind of outside the box, right? Another world.

Before 3.5 billion years ago, probably the form of photosynthesis was the primitive anoxygenic form. Why? 

Because aqueous Fe2 was its electron donor, not water. And because we calculated very low oxygen contents, something like these purple bacteria would survive today, but in very restricted environments. By about 3.2 billion years old, we think more advanced oxygenic photosynthesis comes. If you can reach the cellular complexity of oxygenic photosynthesis, you've got the world at your feet. Because your requirements are just water and CO2. It's a complex cellular machinery to get that to fix CO2 to organic carbon and to break apart oxygen, or water to make oxygen. But once you make that, you've got an unlimited energy source, essentially, as long as the sun is shining. But this oxygen at 3.2 billion years ago was only in the shallow oceans because those are the deposits that we analyzed, right? These deposits were deposited in the ocean basins. So those cyanobacteria, good case. 

We don't know that that's exactly what they were, but something like modern cyanobacteria were probably flourishing by about 3.2 billion years ago. The products of this cyanobacteria is to make this iron 3 oxides and a bunch of organic carbon. They die in the shallow oceans, and they settle down to the bottom of the ocean floor. Ah! That's the fuel for another metabolism, the iron reducers, like these guys. And they pump electrons by these little pilis to those iron oxides. They reduce it, and that's how they form energy. And so we saw that first part of that negative iron isotope signal. That's where we kind of say that that's where iron reduction occurred. Maybe it was across the bacteria in the archaea. 

Maybe that puts a pinning point. Aw, that's a little bit more debatable, but it's certainly what we would suggest. Five hundred million years later, the peak of iron formation deposition was there. We were raining out lots of iron oxides, lots of rust, lots of organic carbon from oxygenic photosynthesis. The iron reducers were going like crazy. And that's where most of our metal comes from. The actual rise of oxygen was 200 million years later, about 2.3 billion. But we had to oxidize all that iron before we would have free oxygen actually accumulate in the atmosphere. Those chemical reactions are important. And the evidence I pointed out would be things like these rusted lake deposits of 2.3 billion years old. 

That was a problem for the iron reducers, because now we had weathering on the continents that dumped out all this thing that we call sulfate, oxidized pyrite. And another metabolism took over microbial sulfate reduction, and the iron reducers lost out. Clear evolutionary change in terms of the order of metabolisms. Now the last thing I just want to mention is, let's fast forward a little bit. That rise of oxygen 2.3 billion years ago was the first one. The rise of oxygen to our current levels didn't happen for another billion or so years, at about 600 million or .6 billion years ago. That was a second rise of oxygen up to levels that we actually needed to support aerobic respiration, what we do. The energy yield of that is high. But until oxygen levels were high enough to support aerobic respiration, animal life could not emerge. And so that's actually relatively late in the planet. 

Now this becomes really important when we say okay, if Earth is our model and we want to go look for life elsewhere, do we expect to find a whole bunch of animal fossils on something like Mars? No. We expect to find microbial life, because you can, a planet seems to sit for a long, long time in a microbial stage. And it's only when we change to a very high energy respiration, at high oxygen levels, aerobic respiration, can you actually have relatively complex life. The first life were these soft-bodied, these soft-bodied animals, very rarely preserved. They're not hard-shelled fossils. So those, those fauna or those, those fossils are about 600 million years old. A little bit younger, we get the hard-shelled fossils. And that's things like the trilobite, right? Many of you have seen that; it's a state fossil. 

And that's when things really exploded. And so once oxygen, the second rise up to modern levels occurred, then the biosphere really took off. Thanks. 
(audience applauding)

 We do this every Wednesday night, 50 times a year.