Sunday, February 7, 2016

{5b} Earth's Earliest Climate - By Angela Hessler

I believe unfamiliarity with our planet's life story is at the root of society's inability to grasp serious climate science.  This in turn makes people frighteningly gullibility when it comes to falling for the most pathetic of con jobs that the Republican/libertarian PR Machine keep producing and broadcasting.

Listen to them deny basic Earth observations and geophysical fundamentals, or that favorite, denying the science by attacking messengers in order to ignore the scientific information.   It's like the Inhofes and Kochs and Christies and Lindzens, all them, possess a mind's eye concept of our planet with the depth of a post card.  

No appreciation whatsoever for the complexity of what we have here, or the eons of evolutionary "tinkering" that created this fantastic planet we were born into.  Nor any conception of the massive momentum that goes into our weather systems and global circulation patterns.  All they can see from within their protective bubble is resources to consume, power and money.  That their attitude is infantile and suicidal don't seem to matter.

That's why I started this project , because I wanted to share some of the learning process that's gone into building my own understanding and appreciation of evolution and in turn our climate system.  Admittedly I'm no scholar, but I sure am a student of my Earth and have some valuable information to share.  I challenge you to try and do a better job.  Please!

In this fifth installment I rely on an expert to present an excellent summation of the state of our understanding regarding the evolution of our climate system.  Ironically, after I finally finished working on condensing Dr. Angela Hessler's paper, (not an easy task for such a compactly written report), I started researching getting permission to post this, only to find the following:
TERMS OF USE - You may reproduce this material, without modifications, in print or electronic form for your personal, non-commercial purposes or for non-commercial use in an educational environment.
Well, okay if that's how it's got to be, he says with a smile, I wasn't feeling that good about my trimming anyways.  I did venture to highlight key sentences.  With no further ado, here's Dr. Angela Hessler's informed grand tour of the evolution of our climate system.

Earth's Earliest Climate
By: Angela M. Hessler | now with the Deep Time Institute  
(Chevron Energy Technology Company) © 2011 Nature Education 

Citation: Hessler, A. M. (2011) Earth’s Earliest Climate. Nature Education Knowledge 3(10):24

When we discuss climate change today, we are mostly concerned with how such change will impact our environment and our lives. We look to the past to help understand climate cycles and how our current anthropogenic changes fit into natural change. Even more, we look to the past to help us find solutions.

Perhaps the most compelling reason to understand “deep-time” climate change is in how it relates to the origin and evolution of life on Earth, and possibly beyond. Despite the stark differences between today’s world and that of the Archean, it is clear that at both times, climate has impacted — and been impacted by — life on Earth.

This paper will take you as far back in the climate record as is currently possible, to the Archean Eon, from 3.9 to 2.5 billion years ago (Bya) (Figure 1). Peering so deeply back in time, far beyond the resolution of many isotope analysis methods, we invariably lose the details about climate and atmosphere chemistry that we can achieve — for instance, analyzing 500,000 year-old gas bubbles in Antarctic ice cores. Instead, we must ask fundamental questions: What was Earth’s surface like? Was its climate hot? Was it icy? Was there a greenhouse effect? For answers, we look to three far-flung Archean terranes.

Isua in West Greenland, Barberton in South Africa, and Pilbara in Western Australia.

{Our Earth's vital signs over the past 4.5 billion years.}

Thin red line, that drops dramatically and then has a couple spikes = Impact rate
Doted red line = Solar luminosity
Dark green line = CO2 
Light green line = CH4 (methane)

Red line = O2
Blue line = Ocean surface temperature
Two vertical bands, in the Archean = earliest sedimentary rock formation (continent building)
Dark gray vertical bands = glacial episodes

(The moon would have formed shortly after accretion ended (within the white bar), early in the Hadean)

(The reign of humanity would fit within the tiniest vertical sliver at the far right end of this graph)

Figure 1: Earth’s changing atmosphere and climate through geologic time.
This diagram combines data and models from select published studies to provide a broad overview of how atmosphere, climate, and life have interacted and changed over the whole of Earth’s history. 
Curves for CH4, CO2, and O2 are based on the PAL (present atmospheric level) and use the logarithmic scale on the right-hand y-axis. 
Ocean surface temperature (°C) uses the right axis. 
Solar luminosity is presented as a percentage of today’s value, and uses the left axis. 
Impact rate is also based on recent rates and uses the left-axis logarithmic scale. 
The source for each curve is indicated by its reference number. 
Dotted lines are areas of conjecture. 
The level of detail for change during the Phanerozoic is largely (but not entirely) a function of geologic preservation and data availability.

A Habitable Young Earth?

A non-hostile climate was prerequisite for the evolution of life as we know it. Was there liquid water? Was there crust on which life could take hold? Earth’s habitable surface today is certainly different than it was 4.6 Bya when it first condensed out of our Sun’s dusty, rotating nebula. We have very little remaining evidence of Earth’s crust during its first 500 million years — just a handful of hardy zircon grains from the Jack Hills conglomerate in western Australia (Maas et al. 1992, Wilde et al. 2001). Despite their tiny size, locked within each zircon’s crystal lattice is a remarkable record of events back 4.4 Bya. Their chemistry, detected by ion beams just a few microns wide, suggest they formed as part of buoyant crust and in the presence of liquid water (Mojzsis et al. 2001, Peck et al. 2001, Wilde et al. 2001). 

If liquid water and elevated crust existed so early in Earth’s history, was life present as well? There is carbon isotope evidence for life in the world’s oldest known volcanic-sedimentary rocks (3.7–3.9 Bya) in the Isua terrane of West Greenland (Rosing 1999). Sedimentary rocks deposited in deepwater, below the photic (light-penetrated) zone, contain muddy, carbon-rich layers with a carbon-isotope signature similar to organic-rich muds of the modern ocean. As in the modern ocean, these carbon-rich layers could have accumulated as planktonic bacteria at the surface died and settled to the ocean floor. Not only had life evolved by 3.7 Bya, but the evidence from Greenland hints that these life forms could have been photosynthetic, thriving in light-filled surface waters.

Life before 3.9 Bya is more conjectural, because older rocks were obliterated by the late heavy meteorite bombardment (LHB) that also cratered our Moon. It may be that extremophiles, which we now understand to live at temperatures up to 120°C (Kashefi & Lovley 2003) and depths more than 3 km (Lin et al. 2006), existed before 3.9 Bya and were suited to survive the LHB by colonizing deep habitats (Abramov & Mojzsis 2009). On the other hand, there is ample evidence that tidal-flat ecosystems survived a period of heavy meteorite bombardment between 3.5–3.2 Bya — impacts 10 to 100 times bigger than the Cretaceous-Tertiary (K-T) event that generated forceful tsunamis (Byerly et al. 2002). It appears unnecessary to assume that life “waited” to evolve until after the LHB events, nor that extremophiles were the only forms suited to survive periods of heavy impacts.

Whether life arose before or after the LHB, it did so without the help of free oxygen. It is well established from sedimentary rocks and paleosols that free oxygen did not accumulate in the atmosphere until after 2.5 Bya (Rasmussen & Buick 1999, Farquhar et al. 2000, Pavlov & Kasting 2002, Holland 2006). In fact, the final rise to modern atmospheric levels occurred only ~580 million years ago (Mya), allowing complex life to diversify on land (Des Marais et al. 1992, Knoll 1992, Canfield & Teske 1996, Narbonne & Gehling 2003).

The Faint Young Sun “Paradox”

There was a shadow over those early years — a much dimmer Sun. At 70% current luminosity (Gough 1981), temperatures at Earth’s surface should have been well below freezing (Sagan & Mullen 1972, Kasting 1993). But in fact the first known widespread glaciation (Marmo & Ojakangas 1984, Evans et al. 1997) did not occur until 2.3 Bya, and by that time, solar luminosity had risen to 83% of present-day levels (Gough 1981). So by 3.9 Bya in West Greenland, where liquid water clearly existed and possibly sustained ancient life, what prevented the Earth from freezing under its “cool” Sun? 

Most research has looked to something like our modern greenhouse effect as a possible solution. Computer models (Kasting 1987, Pavlov et al. 2000, 2003, Catling et al. 2001, Goldblatt et al. 2006, Haqq-Misra et al. 2008, Rosing et al. 2010) offer suggestions as to how this greenhouse may have looked, but they need the ground-truth provided by geologic evidence. Greenland’s Isua terrane contains the world’s oldest sedimentary rocks — however, these ancient sediments were deposited in deep water and did not come in direct contact with Earth’s atmosphere. The oldest known subaerial sediments — those deposited on land in direct contact with the atmosphere — are preserved in the Barberton terrane (South Africa) and the Pilbara terrane (Australia).

Our Earliest Climate Record

From the Isua terrane at 3.7 Bya, we can fast-forward about 500 million years to South Africa, where the oldest preserved emergent crust — still outcropping today in the Barberton Greenstone Belt, near Swaziland — was splashed with rain, traversed by rivers, and colonized by microbial mats more than 3.2 Bya. Together with the Pilbara block in Western Australia, these rocks are our oldest subaerial deposits, our oldest direct evidence of interactions between the atmosphere and geosphere. The rocks are, therefore, our oldest record of climate. 

What do the Barberton rocks tell us about surface temperature? First, there is no evidence of deposition or erosion by glaciers — no poorly-sorted till, dropstones, or glacial striations on bedrock. Second, there is ample evidence of deposition by liquid water — wave ripples in sandstone, sandstone-shale couplets deposited by tides, dune crossbeds built by storm waves, and well-rounded cobbles transported by rivers (Figure 3). From this evidence of freely moving liquid water, we can safely presume that surface temperatures were somewhere between 0–100°C. We already know there was freely moving water back at 3.7 Bya from the marine sedimentary rocks of the Isua terrane. So what new information do the Barberton and Pilbara terranes offer us?
Figure 3: Surface processes preserved in the Barberton Greenstone Belt, South Africa.
(a) thin sandstone-shale laminations from shallow marine deposits; (b) current ripples in sandstone from river deposits (arrow indicates direction of flow); (c) dune-scale crossbedding from shallow marine deposits; (d) core from a gold mine showing well-rounded cobbles transported by rivers.

To better pinpoint temperature at 3.2–3.5 Bya, several studies (Knauth & Lowe 2003, Hren et al. 2009, Blake et al. 2010) have measured stable oxygen isotope (δ18O) values preserved in Barberton rocks. The most common rock studied here is chert, a rock whose silica (1) precipitated directly from fluid (i.e., seawater, hydrothermal fluids), (2) accumulated as the silica-rich shells of marine organisms on the seafloor, or (3) formed as the result of silicification of other “precursor” minerals. Chert is abundant in the Barberton terrane, and because chert can precipitate directly from seawater, these rock samples are thought to preserve valuable information about the ancient oceans. 

In particular, chert contains stable oxygen isotopes, which (along with stable carbon isotopes) are the basis for estimating ocean temperatures over recent geologic time (Zachos et al. 2007). For recent geologic time, these estimates are usually based on rocks that contain carbonate (CaCO3, i.e., limestone). But carbonate rock is not preserved in Barberton, while chert is abundant. In the case of chert, here’s how the isotope-temperature relationship works: as chert precipitates from seawater, it captures variable amounts of the isotope 18O, depending on the temperature (T) of the seawater. The relationship is mathematically defined (Clayton et al. 1972): (3.09 x 106 T-2) – 3.29 = δ18Ochert – δ18Oseawater. For carbonate rocks, the precipitation reactions are different and so is the equation.

The Barberton cherts were found to be strongly depleted in 18O, giving ocean temperature estimates of ~70°C, more than four times the average temperature of our modern ocean. These results are based on the assumption that oxygen isotope values in the Archean ocean were similar to those in the modern ocean, which remains a point of contention (Kasting et al. 2006, Jaffres et al. 2007). Recently, researchers revisited the Barberton chert using the combined relationship of two stable isotoptes31: δ18O and δD (hydrogen); they found the chert to have precipitated from ocean water no warmer than 40°C. A subsequent study looking at the 18O of Barberton phosphates (Blake et al. 2010) placed Archean ocean temperatures at 26–35°C, corroborating the lower temperatures (Hren et al. 2009) but also supporting the assumption that isotope values of the ocean have not evolved (Knauth & Lowe 2003)

Two additional observations from the Barberton rocks lean toward a more clement climate during the Archean. First, at the proposed temperature of 70°C (Knauth & Lowe 2003), we would expect to see evidence of readily dissolved silica (Sleep & Hessler 2006), just as we expect to see rapid dissolution of limestone in hot, humid climates today. This could manifest at a variety of scales: (1) microscopic etching on the surface of quartz grains in sandstone; (2) micro-karst features on chert cobbles in conglomerates; and (3) karst topography preserved on the tops of chert beds. To date, such features are not observed in Archean rocks, which suggests (but does not require) that surface temperatures were below ~70°C. 

Secondly, the detritus in Barberton sedimentary rocks underwent degree of chemical weathering between that occurring today in warm-temperate and tropical environments (Hessler & Lowe 2006). It is tempting, then, to conclude that average temperature during the Archean was in the range of 18–24°C, like today’s temperate to tropical climes. However, other variables are at play during weathering: the amount of rainfall, the acidity of weathering fluids, and the length of time the detritus was exposed to weathering. For instance, temperatures above 24°C may have been necessary to achieve a high degree of weathering if detritus were exposed to weathering for only a short time. In the Archean, long before rooted land plants were around to stabilize soils, this may have been the case.

An Ancient Greenhouse

Evidence points to an unfrozen — perhaps balmy — Archean Earth, despite a faint Sun. Was a greenhouse atmosphere responsible? To warm the planet above freezing, models show that the Archean atmosphere would have needed 100–1000 times more CO2 than present atmospheric level (PAL) (Kasting 1993). Again, we can look to the sedimentary rocks of the Barberton terrane for hard evidence of the ancient concentration of atmospheric CO2. This is not as straightforward as it is for our recent climate record (<500,000 years), where the atmosphere is literally preserved as bubbles in glacial ice. But there are minerals within the Barberton rocks that require certain amounts of CO2 in order to precipitate: siderite (FeCO3) and nahcolite (NaHCO3). Example equilibrium reactions show that, depending on temperature, increasing carbon dioxide on the right will force the reaction to the left, to precipitate either siderite (Fe-system) or nahcolite (NaH-system). 

3 FeCO3 + 2 SiO2 + 2 H2O <—> Fe3Si2O5(OH)4 +3 CO2 
2 NaHCO3 <—> Na2CO3 + H2O + CO2 

In the Barberton terrane, siderite is preserved within weathering rinds on river cobbles in 3.2 Bya conglomerates (Hessler et al. 2004), and nahcolite is preserved within 3.4 Bya marine sedimentary rocks (Lowe & Tice 2004). At 25°C, the minimum amount of CO2 required to precipitate these minerals is in the range of 7–10 PAL (Eugster 1966). This range is well below that required by the models (Kasting 1993), but these are also minimum values and by themselves do not preclude higher CO2 . A more recent study (Rosing et al. 2010) gives an upper limit for Archean CO2 to be within 10 PAL, based on the presence of magnetite (Fe3O4) in equilibrium with siderite in Archean shallow-marine sediments; higher CO2 levels would keep Fe2+ in solution in the ocean, preventing precipitation of magnetite. All geologic evidence taken together, the Archean atmosphere likely had <10 PAL of CO2 , not enough greenhouse to counteract the cooler Sun, based on earlier models (Kasting 1993). 

If the concentration of atmospheric CO2 in the Archean was insufficient, perhaps another greenhouse gas helped fill the gap. Methane (CH4) seems a good candidate. Archean organisms may have been prolific CH4producers, prior to the advent of oxygen photosynthesis (Kharecha et al. 2005). An anoxic ecosystem — photosynthetic bacteria producing organic matter, and methanogenic archea consuming organic matter and producing methane (Tice & Lowe 2006)—thrived in shallow marine to tidal flat environments in the Barberton terrane (Noffke et al. 2006, Heubeck 2009), 3.2–3.4 Bya and in the Pilbara terrane (Lowe 1983, Hofmann et al. 1999, Allwood et al. 2006). And in the low-oxygen Archean atmosphere, CH4 would have had a longer residence time (Zahnle 1986, Pavlov et al. 2000) to build up to higher concentrations than it does today. 

However, laboratory experiments have shown that when atmospheric CH4 concentration approaches that of CO2, a hydrocarbon haze is produced (Trainer et al. 2006): a real example is Saturn’s moon, Titan. Because a hydrocarbon haze blocks sunlight, a CH4–rich atmosphere runs the risk of self-cooling. In other words, an “anti-greenhouse” would be created, sending Archean temperatures back toward freezing. Therefore, it is likely that while a modest CO2 - CH4 greenhouse existed (and could have included a thin hydrocarbon haze), a high concentration CO2 - CH4 greenhouse may not be plausible (Haqq-Misra et al. 2008). 

Recent models point to increased nitrogen (Goldblatt et al. 2006) as having the ability to enhance a CO2- CH4 greenhouse, because N2 increases atmospheric pressure, in turn increasing the amount of radiation that CO2and CH4 can absorb. Doubling N2 (PAL) could lead to a 4.4°C temperature increase. Where did all this N2 come from and how was it later removed from the atmosphere to reach current levels? Researchers (Mather et al. 2004, Goldblatt et al. 2006) propose that volcanic outgassing (more prevalent during the Archean due to higher heat flow) contributed N2 (some already “fixed” (Mather et al. 2004) as volcanic NOx and/or NH3) to the atmosphere early on. After the origin of photosynthesis, N2 was removed from the atmosphere and sequestered in rocks through nitrogen-fixation (N2 + 3 H2O --> 2 NH3 + 1.5 O2), the removal rate dramatically increasing once the deep oceans turned oxic in the late Proterozoic. Nitrogen as part of Earth’s early atmosphere is particularly relevant to the origin of life — its fixed form is considered necessary for biological synthesis.

Up in the Clouds

Recent studies have suggested that lower planetary albedo helped heat the early Earth (Ronandelli & Lindzen 2010, Rosing et al. 2010). Albedo is a measure of Earth’s reflectance — higher albedo means that the Earth is reflecting more sunlight back to space, rather than trapping it as heat. Earth’s albedo varies with different surfaces: for example, deserts are highly reflective, while forested land and oceans are not. Albedo during the Archean may have been lower than today if we assume the following (Rosing et al. 2010): (1) continental area was smaller and thus less reflective, and (2) clouds were more transparent and thus less reflective. We already discussed that some “buoyant” (continental-like) crust may have existed as early as 4.4 Bya, based on zircon grains from the Australian Jack Hills. However, the oldest preserved, clearly emergent blocks (3.2–3.5 Bya) occur only in the Barberton and Pilbara regions, and continents appear to have gained mass ever since. 

Why would Archean clouds be different, more transparent, than they are today? Life, once again, may have been a player. Today, clouds are effective at reflecting sunlight back to space due to the abundance of cloud condensation nuclei (CCN), solid surfaces onto which water vapor condenses and forms water droplets. Non-organic particles (dust, salt, soot, etc.) serve as cloud condensation nuclei. Over the oceans, 50% of these CCN are formed by the oxidation of sulfuric gases released to the atmosphere by plants and eukaryotic algae (Kreidenweis & Seinfeld 1988). A recent model proposes (Rosing et al. 2010) that in the Archean, before the evolution of eukaryotes, there was essentially no biogenic sulfuric gas available to promote cloud formation, and therefore clouds were not as dense and reflective as they are today. 

The concept of cloud albedo having a “stabilizing effect” on Earth’s early climate is long held (Rossow et al. 1982). Renewed interest (Rondanelli & Lindzen 2010, Rosing et al. 2010) in the relative importance of albedo versus greenhouse gases has come about because, as discussed earlier, geologic evidence does not support the required high levels of standard greenhouse gases (CH4, CO2) in counteracting the “faint young Sun”. Still, the relationship between clouds, life, and climate is a complex one, and there is considerable uncertainty in applying this process to the early Earth (Goldblatt & Zahnle 2011).

Oxygen and Glaciers: A More Familiar World

From the 3.2 Bya river deposits of the Barberton terrane, fast-forward another 500 million years, to the close of the Archean Eon, now preserved in the Pilbara region of western Australia. It was thought that the oldest known evidence of cyanobacteria was preserved in hydrocarbon “biomarkers” in 2.7 Bya Pilbara rocks (Brocks et al. 1999, Summons et al. 1999), but these biomarkers have since been refuted as younger contaminants (Rasmussen et al. 2008) and as non-unique to oxygen-producing bacteria (Rashby et al. 2007). The evolution of oxygenic photosynthesis may have begun prior to 2.5 Bya, based on evidence for weak oxidation of minerals in shales of the Pilbara terrane (Anbar et al. 2007). However, the real turnover took place at 2.45 Bya. The “Great Oxidation Event” was complete by 2.32 Bya (Bekker et al. 2004), quickly transforming the Earth’s atmosphere from essentially anoxic to its near-present state. Following this major transition, oxidized minerals were widespread in sedimentary rocks around the world (Rasmussen & Buick 1999, Farquhar et al. 2000, Pavlov & Kasting 2002, Holland 2006), and the first known cyanobacteria (Hofmann 1976) and eukaryotic (Knoll et al. 2006) fossils appear in the rock record. Organisms had “learned” oxygen photosynthesis and metabolism, and Earth has not since returned to an anoxic atmosphere. 

At the same time, Earth experienced its first known interval of global glaciations (Marmo & Ojakangas 1984, Evans et al. 1997). Did oxygen-exhaling organisms cause Earth’s first Ice Age? Perhaps, because methane would have been removed by oxidation from the atmosphere, converted to the less-effective greenhouse gas CO2 (Kopp et al. 2005, Bekker & Kaufman 2007). The greenhouse may have been additionally weakened by other events just preceding the 2.4 Bya glaciation: major continental rifting events that forced atmospheric CO2 to be absorbed during the weathering of newly exposed crust (Evans 2003, Melezhik 2006) (i.e., CaSiO3 + CO2 + H2O —> Ca2+ + HCO3- + H4SiO4), and subsequently, the precipitation of abundant marine carbonate (i.e., Ca2+ + HCO3- —> CaCO3 + H+) in newly-opened tropical seaways.

Whatever the cause-and-effect, the early Proterozoic glaciation marked the rise of oxygen-producing cyanobacteria to build our modern oxygenic atmosphere, and the subsequent decline of methane-producing microbes and their carbon-rich atmosphere. Another major step occurred less than 5 million years prior to — and probably prompting (Knoll 1992) — the “Cambrian explosion” of complex life onto land (Narbonne & Gehling 2003), with O2 making a final ascent to modern levels in both the atmosphere and deep ocean around 580 Mya (Canfield et al. 2006). Methane potentially declined in response. With a weaker greenhouse (O2 mostly replacing CO2, CH4), and with abundant eukaryotic algae to generate CCN and more reflective clouds, climate for the remainder of Earth’s history has hovered (despite a steadily brightening Sun) more narrowly above freezing — nudged into “icehouse” and “greenhouse” conditions with perhaps smaller perturbations in the greenhouse-albedo system.

Deep-Time Climate Change

We currently inhabit an “icehouse” world. We use much of the available data and geologic evidence to understand near-term (<1 million years) climate fluctuations typical of icehouse periods and what this might mean for our future. While the “greenhouse” worlds of the Archean, Proterozoic, and much of the Phanerozoic may seem foreign in many ways, these times do reveal Earth’s full spectrum of known climate possibilities. 

Perhaps what is most telling is that climate regulation has been a mainstay from the beginning. Despite large changes in solar energy as well as dramatic impact events, our climate has been perpetually suitable for some form of life. Inorganic processes have played a big part in this regulation, particularly through cycles of outgassing, weathering, albedo, and oceanic circulation associated with plate tectonics. 

From its origin, life has greatly impacted its climate-atmosphere system — without permanently tipping the balance toward uninhabitability. However we also see that when conditions reach a tipping point (e.g., The Great Oxidation Event), change can be extraordinarily rapid and (as yet) irreversible. As we continue to impact the ocean-atmosphere system, we must look to deep-time climate change — particularly these abrupt and seemingly permanent transitions — to more fully frame our forecasts and design our solutions.

Extremophile: An organism that thrives in physically or geochemically extreme conditions, including but not limited to: high temperature (> 60° C), low temperature (< -15° C), high pressure, high acidity (pH < 3), high alkalinity (pH > 9), high salinity (> 0.2 M NaCl), low humidity, low nutrient availability, low oxygen, and/or high levels of radioactivity
Paleosol: A former or "fossil" soil that is preserved within sedimentary or volcanic deposits

References and Recommended Reading

Abramov, O. & Mojzsis, S. J. Microbial habitability of the Hadean Earth during the late heavy bombardment. Nature 459, 419–422 (2009).

Allwood, A. C. et al. Stromatolite reef from the early Archean era of Australia. Nature 441, 714–718 (2006).

Anbar, A. D. et al. A whiff of oxygen before the Great Oxidation Event? Science 317, 1903–1906 (2007).

Beerling, D. et al. Methane and the CH4-related greenhouse effect over the past 400 million years. American Journal of Science 309, 97–113 (2009).

Bekker, A. & Kaufman, A. J. Oxidative forcing of global climate change; A biogeochemical record across the oldest Paleoproterozoic ice age in North America. Earth and Planetary Science Letters 258, 486–499 (2007).

Bekker, A. et al. Dating the rise of atmospheric oxygen. Nature 427, 117–120 (2004).

Berner, R. A. GEOCARBSULF: A combined model for Phanerozoic atmospheric O2 and CO2. Geochimica et Cosmochimica Acta 70, 5653–5664 (2006).

Berner, R. A. Phanerozoic atmospheric oxygen: New results using the GEOCARBSULF model. American Journal of Science 309, 603–606 (2009).

Blake, R. E., Chang, S. J. & Lepland, A. Phosphate oxygen isotope evidence for a temperate and biologically active Archean ocean. Nature 464, 1029–1033.

Brocks, J. J. et al. Archean molecular fossils and the early rise of Eukaryotes. Science 285, 1033–1036 (1999).

Byerly, G. R. et al. An Archean impact layer from the Pilbara and Kaapvaal cratons. Science 297, 1325–1327 (2002).

Canfield, D. E. & Teske, A. Late Proterozoic rise in atmospheric oxygen concentration inferred from phylogenetic and sulphur-isotope studies. Nature 382, 127–132 (1996).

Canfield, D. E., Poulton, S. W. & Narbonne, G. M. Late-Neoproterozoic deep-ocean oxygenation and the rise of animal life. Science 315, 92–95 (2006).

Catling, D. C., Zahnle, K. J. & McKay, C. P. Biogenic methane, hydrogen escape, and the irreversible oxidation of early Earth. Science 293, 839–843 (2001).

Clayton, R. N., O’Neil, J. R. & Mayeda, T. K. Oxygen isotope exchange between quartz and water. Journal of Geophysical Research 77, 3057–3067 (1972).

Des Marais, D. J. et al. Carbon isotope evidence for the stepwise oxidation of the Proterozoic environment. Nature 359, 605–609 (1992).

Eugster, H. P. Sodium carbonate-bicarbonate minerals as indicators of PCO2. Journal of Geophysical Research 71, 3369–3378 (1966).

Evans, D. A. A fundamental Precambrian–Phanerozoic shift in Earth’s glacial style? Tectonophysics 375, 353–385 (2003).

Evans, D. A., Beukes, N. J. & Kirschvink, J. L. Low-latitude glaciations in the Paleoproterozoic era. Nature 386, 262–266 (1997).

Farquhar, J., Bao, H. & Thiemans, M. Atmospheric influences of Earth’s earliest sulfur cycle. Science 289, 756–758 (2000).

Goldblatt, C. & Zahnle, K. J. Clouds and the faint young Sun paradox. Climate of the Past 7, 203–220 (2011).

Goldblatt, C., Lenton, T. M. & Watson, A. J. Bistability of atmospheric oxygen and the Great Oxidation. Nature 443, 683–686 (2006).

Gough, D. O. Solar interior structure and luminosity variations. Solar Physics 74, 21–34 (1981).

Haqq-Misra, J. D. et al. Revised, hazy methane greenhouse for the Archean Earth. Astrobiology 8, 1127–1137 (2008).

Hessler, A. M. & Lowe, D. R. Weathering and sediment generation in the Archean: An integrated study of the evolution of siliciclastic sedimentary rocks of the 3.2 Ga Moodies Group, Barberton Greenstone Belt, South Africa. Precambrian Research 151, 185–210 (2006).

Hessler, A. M. et al. A lower limit for atmospheric carbon dioxide levels 3.2 billion years ago. Nature 428, 736–738 (2004).

Heubeck, C. An early ecosystem of Archean tidal microbial mats (Moodies Group, South Africa, ca. 3.2 Ga). Geology 37, 931–934 (2009).

Hofmann, H. J. Precambrian microflora, Belcher Islands, Canada: Significance and systematics. Journal of Paleontology 50, 1040–1073 (1976).

Hofmann, H. J. et al. Origin of 3.45 Ga coniform stromatolites in Warrawoona Group, Western Australia. Geological Society of America Bulletin 111, 1256–1262 (1999).

Holland, H. D. The oxygenation of the atmosphere and oceans. Philosophical Transactions of the Royal Society B: Biological Sciences 361, 903–915 (2006).

Hren, M. T., Tice, M. M. & Chamberlain, C. P. Oxygen and hydrogen isotope evidence for a temperate climate 3.42 billion years ago. Nature 205, 205–208 (2009).

Jaffres, J. B. D., Shields, G. A. & Wallmann, K. The oxygen isotope evolution of sea water; A critical review of a long standing controversy and an improved geological water cycle model for the past 3.4 billion years. Earth-Science Reviews 83, 83–122 (2007).

Kashefi, K. & Lovley, D. R. Extending the upper temperature limit for life. Science 301, 934 (2003).

Kasting, J. F. Theoretical constraints on oxygen and carbon dioxide concentrations in the Precambrian atmosphere. Precambrian Research 34, 205–229 (1987).

Kasting, J. F. Earth’s early atmosphere. Science 259, 920–926 (1993).

Kasting, J. F., Liu, S. C. & Donahue, T. M. Oxygen levels in the prebiological atmosphere. Journal of Geophysical Research 84, 3097–3107 (1979).

Kasting, J. F. et al. Paleoclimates, ocean depth, and the oxygen isotopic composition of seawater. Earth and Planetary Science Letters 252, 82–93 (2006).

Kharecha, P., Kasting, J. & Seifert, J. A. A coupled atmosphere-ecosystem model of the early Archean Earth. Geobiology 3, 53–76 (2005).

Knauth, L. P. & Lowe, D. R. High Archean climatic temperatures inferred from oxygen isotope geochemistry of cherts in the 3.5 Ga Swaziland Supergroup, South Africa. Geological Society of America Bulletin 155, 566–580 (2003).

Knoll, A. H. The early evolution of eukaryotic organisms: A geological perspective. Science 256, 922–627 (1992).

Knoll, A. H. et al. Eukaryotic organisms in Proterozoic oceans. Philosophical Transactions of the Royal Society B: Biological Sciences 361, 1023–1038 (2006).

Kopp, R. E. et al. The Paleoproterozoic snowball Earth: A climatic disaster triggered by the evolution of oxygenic photosynthesis. Proceedings of the National Academy of Sciences of the United States of America 102, 11131–11136 (2005).

Kreidenweis, S. M. & Seinfeld, J. H. Nucleation of sulfuric acid-water and methanesulfonic acid-water solution particles: Implications for the atmospheric chemistry of organosulfur species. Atmosphere Environment 22, 283–296 (1988).

Lin, L-H. et al. Long-term sustainability of a high-energy, low-diversity crustal biome. Science 314, 479–482 (2006).

Locklair, R. E. & Lerman, A. A model of Phanerozoic cycles of carbon and calcium in the global ocean: Evaluation and constraints on ocean chemistry and input fluxes. Chemical Geology 217, 113–126 (2005).

Lowe, D. R. Restricted shallow water sedimentation of Early Archean stromatolitic and evaporitic strata of the Strelley Pool Chert, Pilbara Block, Western Australia. Precambrian Research 19, 239–283 (1983).

Lowe, D. R. & Tice, M. M. Geologic evidence for Archean atmosphere and climatic evolution: Fluctuating levels of CO2, CH4, and O2 with an overriding tectonic control. Geology 32, 493–496 (2004).

Mather, T. A., Pyle, D. M. & Allen, A. G. Volcanic source for fixed nitrogen in the early Earth’s atmosphere. Geology 32, 905–908 (2004).

Maas, R. et al. The Earth’s oldest known crust: A geochronological and geochemical study of 3900–4200 Ma detrital zircons from Mt. Narryer and Jack Hills, Western Australia. Geochimica et Cosmochimica Acta 56, 1281–1300 (1992).

Marmo, J. S. & Ojakangas, R. W. Lower Proterozoic glaciogenic deposits, eastern Finland. Geological Society of America Bulletin 98, 1055–1062 (1984).

Melezhik, V. A. Multiple causes of Earth’s earliest global glaciations. Terra Nova 18, 130–137 (2006).

Mojzsis, S. J., Harrison, T. M. & Pidgeon, R. T. Oxygen-isotope evidence from ancient zircons for liquid water at the Earth’s surface 4,300 Myr ago. Nature 409, 178–181 (2001).

Narbonne, G. M. & Gehling, J. G. Life after snowball: The oldest complex Ediacaran fossils. Geology 31, 27–30 (2003).

Noffke, N. et al. A new window into Early Archean life: Microbial mats in Earth’s oldest siliciclastic tidal deposits (3.2 Ga Moodies Group, South Africa). Geology 34, 253–256 (2006).

Pavlov, A. A. & Kasting, J. F. Mass-independent fractionation of sulfur isotopes in Archean sediments: Strong evidence for an anoxic Archean atmosphere. Astrobiology 2, 27–41 (2002).

Pavlov, A. A. et al. Greenhouse warming by CH4 in the atmosphere of early Earth. Journal of Geophysical Research 105, 11981–11990 (2000).

Pavlov, A. A. et al. Methane-rich Proterozoic atmosphere? Geology 31, 87–90 (2003).

Peck, W. H. et al. Oxygen isotope ratios and rare earth elements in 3.3 to 4.4 Ga zircons: Ion microprobe evidence for high delta O-18 continental crust and oceans in the Early Archean. Geochimica et Cosmochimica Acta 65, 4215–4229 (2001).

Rashby, S. E. et al. Biosynthesis of 2-methylbacteriohopanepolyols by an anoxygenic phototroph. Proceedings of the National Academy of Sciences of the United States of America 104, 15099–15014 (2007).

Rasmussen, B. & Buick, R. Redox state of the Archean atmosphere: Evidence from detrital heavy minerals in ca. 3250–2750 Ma sandstones from the Pilbara Craton, Australia. Geology 27, 115–118 (1999).

Rasmussen, B. et al. Reassessing the first appearance of eukaryotes and cyanobacteria. Nature 455, 1101–1105 (2008).

Rondanelli, R. & Lindzen, R. S. Can thin cirrus clouds in the tropics provide a solution to the faint young Sun paradox? Journal of Geophysical Research 115, 689–690 (2010).

Rosing, M. T. 13C-depleted carbon microparticles in >3700-Ma sea-floor sedimentary rocks from west Greenland. Science 283, 674–676 (1999).

Rosing, M. T. et al. No climate paradox under the faint Sun. Nature 464, 744–747 (2010).

Rossow, W. B., Henderson-Sellers, A. & Weinreich, S. K. Cloud feedback: A stabilizing effect for the early Earth? Science 217, 1247–1247 (1982).

Sagan, C. & Mullen, G. Earth and Mars – Evolution of atmospheres and surface temperatures. Science 177, 52–56 (1972).

Sheldon, N. D. Precambrian paleosols and atmospheric CO2 levels. Precambrian Research 147, 148–155 (2006).

Sleep, N. H. & Hessler, A. M. Weathering of quartz as an Archean climatic indicator. Earth and Planetary Science Letters 241, 594–602 (2006).

Summons, R. E. et al. 2-methyl-hopanoids as biomarkers for cyanobacterial oxygenic photosynthesis. Nature 400, 554–557 (1999).

Tice, M. M. & Lowe, D. R. Hydrogen-based carbon fixation in the earliest known photosynthetic organisms. Geology 34, 37–40 (2006).

Trainer, M. G. et al. Organic haze on Titan and the early Earth. Proceedings of the National Academy of Sciences of the United States of America 103, 18035–18042 (2006).

Valley, J. W. et al. A cool early Earth. Geology 30, 351–354 (2002).

Wilde, S. A. et al. Evidence from detrital zircons for the existence of continental crust and oceans on the Earth 4.4 Gyr ago. Nature 409, 175–178 (2001).

Zachos, J. et al. Trends, rhythms, and aberrations in global climate 65 Ma to present. Science 292, 686–693 (2001).

Zahnle, K. J. Photochemistry of methane and formation of hydrocyanic acid (HCN) in the Earth’s early atmosphere. Journal of Geophysical Research 91, 2819–2834 (1986).

Johnston, D. T. et al. Anoxygenic photosynthesis modulated Proterozoic oxygen and sustained Earth’s middle age. Proceedings of the National Academy of Sciences of the United States of America 106, 16925–16929 (2009).

Wednesday, January 6, 2016
{1} Our Global Heat and Moisture Distribution Engine

Saturday, January 9, 2016
{2} Co-evolution of Minerals and Life | Dr Robert Hazen

Thursday, January 14, 2016
{3} Evolution of Carbon and our biosphere - Professor Hazen focuses on the element Carbon

Saturday, January 23, 2016
{4} Evolution-Considering Deep Time and a Couple Big Breaks

Saturday, February 6, 2016
{5a} The Most Beautiful Graph on Earth - A. Hessler

Sunday, February 7, 2016
{5b} Earth's Earliest Climate - By Angela Hessler

Sunday, February 14, 2016
{6} Evolution of Earth's Atmosphere - easy version

Thursday, February 18, 2016
{7} Our Global Heat and Moisture Distribution Engine, visualized

Friday, February 19, 2016
{8} Atmospheric Insulation Explained - appreciating our climate engine

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