Web Supplement 25.4

25.4 PAST VARIATIONS IN CLIMATE AND ATMOSPHERIC CARBON DIOXIDE LEVELS

The geologic carbon cycle was described in Chapter 25.1 as providing the negative feedback loop responsible for stabilizing P over time scales of 105 to 107 y. In CO2 this cycle, the burial of CaCO3 and SiO2 in ocean sediments fuels decarbonation reactions that produce CO2 and CaSiO3. The CO2 is degassed through volcanoes and the CaSiO3 is uplifted onto land. Weathering of uplifted CaSiO3 by atmospheric CO2 supplies bicarbonate, silicate, and calcium ions to river runoff. In the modern ocean, plankton transform these ions back into CaCO3 and SiO2. If the rate of weathering is limited by P , increasing P should lead to enhanced uptake rates as a result of enhanced CO2 CO2 weathering, thereby providing a negative feedback that acts to stabilize P and the CO2 sizes of the other crustal carbon reservoirs. Nevertheless, large transient swings in P CO2 have occurred during various periods of Earth’s history. A few have led to long-term reorganizations of the global carbon cycle. Paleoceanographers are particularly inter- ested in studying these events as they provide some clues as to what changes we can expect as a result of our anthropogenic perturbations. A key to understanding these past swings is recognizing the important role of tectonism in the geologic carbon cycle. First, it determines the rates and loca- tions of the decarbonation reactions and of uplift. Second, it determines the spatial extent of shallow shelves. The latter was critically important prior during the Pre- cambrian and first half of the Phanerozoic because carbonate burial was restricted to the shallow shelves. In the mid-Phanerozoic, the evolution of pelagic calcifiers enabled burial of carbonate in deep-sea sediments. As discussed in Chapter 15.6, deep-sea sedimentary carbonate participates in a stabilizing feedback loop called calcite compensation that operates over time-scales of 104 y. In the present-day ocean, sedimentary carbonate deposition is concentrated in the Atlantic and Indian Oceans (Figure 15.5), whereas subduction, and, hence, decarbonation, is occurring primarily in the Pacific Ocean. Thus, the stability afforded by the geologic carbon cycle will occur only over time scales long enough to capture a supercontinent cycle. The stabilizing feedback of the geologic carbon cycle is also predicted to be of minimal 1 2 Web Supplement 25.4

effect during periods of tectonic stability due to the lack of mountain building and subduction. Tectonism is also important in the geological carbon cycle because it determines land-mass distributions. In the present day, most of the land mass is located in northern hemisphere mid-latitudes. A land-mass distribution in which most of the continents are located at tropical latitudes would be expected to result in a cold climate due to enhanced continental weathering rates. Conversely, our present geography should support a relatively warm climate. This was not the case, at least prior to human-induced warming. The probable explanation is that the geological carbon cycle’s stabilizing feedback is not presently working well. As noted earlier, most subduction is occurring in the Paciﬁc where the sediments have a low carbonate content, thus disabling the geologic carbon cycle’s metamorphic decarbonation linkage. On the other hand, the presence of large land masses in the mid-latitudes does provide some climate stability in that an extreme cooling event leads to the formation of continental ice sheets. This slows continental weathering, permitting volcanic CO2 to reaccumulate and warm the atmosphere. In the following subsections, the long-term evolution of the global carbon cycle is discussed along with the most likely explanations for the large transient swings in P CO2 that have occurred sporadically throughout Earth’s history. The marine carbon cycle has played an important role in the past and is likely to determine the degree to which our perturbations of the global carbon cycle will result in global climate change.

25.4.1 The Carbon Cycle in the Precambrian

When viewed over the long term, atmospheric CO2 levels have been in decline since the early Precambrian. As shown in Figure W25.1, P is estimated to have dropped CO2 by a factor of 100 to 10,000 over Earth’s history, indicating a large-scale relocation of carbon in the crustal-ocean-atmosphere factory. During the Hadean (4.6 to 3.8 bybp), the global carbon cycle was controlled solely by geology in which the reactions shown in Eqs. 25.1 through 25.8 stabilized P through the formation of carbonates via abio- CO2 genic precipitation in warm, shallow-water environments. This stabilizing feedback was important as the Sun’s solar luminosity has been increasing (slowly) over time. On the early Earth, solar luminosity was only 70% of the present day, so a long-term increase in insolation should lead to an overall decline in P assuming that warming temperatures CO2 enhance weathering rates. The advent of life around 3.8 bybp added a new crustal reservoir to the global carbon cycle, that of sedimentary organic matter. (Land plants did not evolve until the Phanero- zoic.) The accumulation of organic matter in a sedimentary reservoir contributed to the overall trend of declining P . One of the earliest life forms were the methanogenic CO2 archaeans. These microbes are thought to have converted virtually all of the primordial volcanic H2 in the atmospheric to methane via reaction with atmospheric CO2. Methane is a far more efﬁcient greenhouse gas than CO2 (Table 25.2). The replacement of CO2 with CH4 is thought to have kept the early Earth’s atmosphere warm and its surface ice free. Prior to the evolution of methanogens, P levels were not high enough to keep CO2 Web Supplement 25.4 3

107 104 Ocean-covered earth

Huronian glaciation 106 (5 to 208C) 103 bar) Neoproterozoic m glaciation 105 (5 to 208C) 102 30% Solar flux 8 reduction (0 C) 104 concentration (PAL)

2 10 partial pressure (

Mt Roe palaeosol 2 CO Constraints provided by the 103 CO Ruyang microfossil analyses 1 Terrestrial C3 photosynthesis 102 0.5 1.5 2.5 3.5 4.5 Time before present (Gyr)

FIGURE W25.1

Atmospheric PCO2 relative to present atmosphere level (PAL) in the Archaean and Proterozoic based on analysis of microfossils (symbols). Shaded area was obtained from modeling. Source: After

Kaufman A. J., and S. Xiao (2003). High CO2 levels in the Proterozoic atmosphere estimated from analyses of individual microfossils. Nature 425, 279–283.

Earth ice free as the Sun’s luminosity was still quite low. Thus, the methanogens added an important element to the carbon cycle that acted on global climate. Methane production by the methanogens did not lead to a runaway greenhouse effect, as a negative feedback was established through the interaction of UV radiation with CH . At high P , UV radiation induces the formation of a photochemical haze 4 CH4 that reﬂects insolation. This methane thermostating ended around 3.0 bybp with the advent of oxygenic photosynthesizers. Their production of O2 led to the oxidation of atmospheric CH4. With the loss of this potent greenhouse gas, the planet entered a period (mid-to-late Proterozoic) in which several global glaciation events occurred. These are referred to as “ice-houses” or “Snowball Earths.” As discussed later, several other causative factors contributed to the occurrence of these ice-house conditions. In contrast to the climate variability of the Proterozoic, no Snowball Earths have occurred since the beginning of the Phanerozoic 550 mybp. The switch back to climate stability is attributed to the advent of multicellular lifeforms and, in particular, to the calcite compensation feedback made possible by evolution of biocalcifying marine organisms.

25.4.2 Snowball Earths Many Snowball Earth episodes are known to have occurred during the Proterozoic. They are recorded as lithiﬁed glacial sedimentary deposits, called tillites, accompanied by banded iron formations (oxidized iron) overlain by very thick carbonate layers, 4 Web Supplement 25.4

called cap carbonates.1 These carbonate deposits are typically 3 to 30 m thick and occur on platforms, shelves, and slopes worldwide. The global distribution of these unique sediments suggest that during a Snowball Earth episode, all the land masses and most, if not all, of the oceans were covered by ice. The ﬁrst series of Snowball Earths occurred between 2.2 and 2.45 bybp at the beginning of the Proterozoic (Huronian and Makganene glaciations). The second series took place during the late Proterozoic (the Sturtian glaciation 710 mybp, the Varanger-Marinoan 635 mybp, and the Gaskiers 582 mybp). This latter period is referred to as the Cryogenian. Evidence suggests that the entire oceans froze during the Cryogenian. Life probably survived under thin ice at equatorial latitudes and at hydrothermal vents.

Entering Snowball Conditions Various combinations of four factors are thought to have led to the runaway cooling that produced Snowball Earth conditions in the Proterozoic. Briefly, these factors are (1) high weathering rates, (2) increased volcanic activity leading to the formation of continental flood basalts, (3) the passage of Earth through a giant molecular cloud every 140 my, and (4) the loss of atmospheric methane following the rise of oxygenic photosynthesis. (The last would have served to cause only the first series of Snowball Earths.) Once cooling had started, a positive feedback is postulated to have occurred when the ice sheets reached some critical latitude due to their cumulative high albedo, i.e., the ice sheets reflect insolation back into space, thereby preventing its absorption by the greenhouse gases. High weathering rates are thought to have been caused by three factors: (1) a pre- ponderance of continents in the tropics, where it is hot and wet; (2) the breakup of supercontinents; and (3) the formation of supermountain chains. During the period of the Snowball Earths, little continental area was located at high latitudes. Instead large polar sea-ice caps were present. They reflected solar radiation but did not cover much land area, leaving the continents ice free and susceptible to weathering. Immediately preceding the Cryogenian, around 830 mybp, a supercontinent, called Rodinia, began breaking up. This continued for nearly 200 my. Weathering rates increase when supercontinents break up. When a supercontinent exists, most land area is far from the ocean and therefore very dry. Conversely, when a supercontinent breaks up into small fragments, formerly arid regions become wetter and, hence, weathering rates increase. Around 650 mybp, just prior to the Minoan glaciation, land fragments generated by the Rodinian breakup collided to form the supercontinent Gondwanaland and gave rise to a supermountain chain, which then began to rapidly erode. The formation of this Trans- gondwanan Supermountain Chain also increased the rate of river input of nutrients to the ocean and is thought to have played a role in the subsequent evolution of multicellular organisms and other eukaryotes. All of these types of enhanced weathering lead to a drawdown in P and, hence, contributed to global cooling. CO2

1 The formation of banded ironstones is discussed in Chapter W8.6.1. Web Supplement 25.4 5

At least some of the Snowball Earths seem to have been preceded by massive continental eruptions of basalt lava, called flood basalts. This happened on land masses very close to the equator and, hence, was subject to intense weathering. This type of basaltic rock weathers rapidly, acting as a rich source of Ca2+ and bicarbonate to the ocean. When P levels dropped at the end of each Snowball Earth episode, CO2 precipitation of this Ca2+ and bicarbonate led to the formation of the massive cap carbonates. Finally, cooling events are also thought to be driven by tropospheric dust loading caused by encounters with interstellar giant molecular clouds (GMCs). Our solar system − encounters a dense GMC (>2000 H atoms cm 3)every109 years or so. Average density GMCs are encountered every 108 years, with greatest probability every ∼140 my as this is the periodicity in which the solar system crosses the galactic spiral arms where GMCs are concentrated. The reflection of insolation by this dust is sufficient to cause Snowball Earth conditions, especially as the dust loading occurs too rapidly to be compensated for by the geological carbon cycle stabilizing feedbacks. Based on this periodicity, Earth has likely encountered approximately four high-density GMCs, each of which could have − triggered a Snowball Earth episode, and 15 lower-density (∼1000 H atoms cm 3) GMCs. The latter are thought to be capable of causing moderate ice ages.

Escape from the Snowball Once a Snowball Earth condition began, cooling would have lowered the silicate weathering rate, stabilizing the climate system at the new colder state. Two other factors would have contributed to lowered weathering rates: (1) at low P , the acidity of CO2 rainwater would be reduced, and (2) the coverage of land by ice prevents weathering. Further stability would be provided by the high albedo contributed by the global ice sheets. Nevertheless, these feedbacks would eventually be overcome by the Geo- logical Carbon Cycle once enough time had elapsed for the decarbonation of calcium silicate and the oxidation of organic matter buried in the sediments prior to the onset of the snowball condition. Continuing volcanic activity would have eventually resupplied CO2 to the atmosphere along with iron to the deep sea through hydrothermal emissions. Conditions of lowered sea level are thought to increase the iron-to-sulfur ratio in hydrothermal emissions. Because the ice isolated the ocean waters from the atmosphere for millions of years, the bottom waters likely became anoxic. This promoted elevated Fe2+ concentrations as a result of anoxic conditions and insufficient sulfide for removal via pyrite precipitation. Once volcanically supplied P rose to some critical level, warming must have CO2 ensued. Melting of the global ice cover is thought to have happened quickly due to various positive feedbacks. For example, a rapid reduction in albedo would result from the pooling of meltwater on ice. Climate modeling suggests that the meltdown could have occurred in as little as 2000 y, leading to a variety of cataclysmic effects. First would have been a rapid and large increase in the rate of river runoff. The flow of freshwater into the ocean would have led to an initial density stratification in which low-density oxic meltwater formed a cap over the denser anoxic seawater. The meltwater would also 6 Web Supplement 25.4

have led to a rapid sea-level rise, ﬂooding continental shelves. As the ice retreated, the rock debris and “ﬂour” produced by mechanical ice weathering would have become exposed and commenced weathering. This likely led to an enhanced runoff of Ca2+ and bicarbonate, causing the shallow waters to become supersaturated with respect to calcium carbonate and thereby enabling formation of more cap carbonates. As deepwater circulation was reestablished and the deepwater oxygenated, dissolved Fe2+ was oxidized, leading to the deposition of the banded iron formations. Because reduced 2+ metals, such as Fe , inhibit precipitation of CaCO3, removal of the metals would have facilitated deposition of more cap carbonates providing more carbon drawdown power. Given the slow pace of the geological carbon cycle, the continental weathering 4 6 reactions must have taken at least 10 to 10 y to lower atmospheric CO2 backtoa steady-state level. This suggests that the end of a Snowball Earth episode was marked by a transient period of high P in which ultra-greenhouse conditions were present. CO2 Warming was probably enhanced by the eventual melt of methane clathrates and, hence, the release of this potent greenhouse gas to the atmosphere. The remobilized methane would have been oxidized fairly rapidly to CO2 via reaction with atmospheric O2.

25.4.3 Phanerozoic Innovations to the Global Carbon Cycle

Oscillations in the P of the atmosphere continued into the Phanerozoic as shown in CO2 Figure W25.2. A continuing correlation with global temperatures suggests that CO2 was an important agent in climate control. Several glacial events have occurred during the Phanerozoic, most notably at 300 mybp and throughout the past 1.8 my. All have been associated with CO2 levels below 1000 ppm with one exception, a short-lived glaciation 450 mybp during which P was approximately 5000 ppm. High-resolution reconstruc- CO2 tions of paleotemperatures and atmospheric P have demonstrated a pervasive, tight CO2 correlation. The record for the past 650,000 y has been obtained from analysis of gas trapped in ice cores retrieved from subpolar and polar latitudes (Figure W25.3). Unlike the Snowball Earth episodes, the glaciations of the Phanerozoic have been restricted to the mid- to high latitudes. Three processes seem to have contributed to the lack of Snowball Earths during the Phanerozoic. First, solar luminosity has continued to increase. Second, after the breakup of Rodinia 830 mybp, subsequent supercontinent cycles have redistributed land masses so that they are no longer concentrated at low latitudes. Third, biological innovations such as evolution of biomineralization and land plants have led to enhanced storage of carbon in the crustal reservoirs and development of stabilizing feedbacks. The Snowball Earth episodes are thought to have played an important role in stimu- lating evolution. The most important consequences were the rapid explosion in diversity of metazoans and eukaryotic phytoplankton at the end of the Proterozoic and beginning of the Phanerozoic (Cambrian period). This is attributed to two side effects of the Snowball Earth episodes: mass extinction events and rapid weathering. The former Web Supplement 25.4 7

8000 Models Measurements 30 GEOCARB III 7000 GEOCARB III Royer Compilation 25 COPSE 30 Myr Filter 6000 Rothman 30 Myr filter 20 5000

4000 COPSE 15 3000 10 2000 Carbon dioxide (ppmv) relative to quaternary relative mean value 2

5 CO 1000 P ROTHMAN 0 NPg K J Tr P C D S O Cm 0 g g o g cccc c 0 100 200 300 400 500 Millions of years ago

FIGURE W25.2

Phanerozoic PCO2 levels as estimated from models and climate proxies. Bars on x-axis denote climatic conditions: glacial (g), cool (c), and oscillating glacial/interglacial (o). For data sources see e Royer, D. L., R. A. Berner, I. P. Montanez, N. J. Tabor, and D. J. Beerling (2004). CO2 as a primary driver of Phanerozoic climate, GSA Today 14, 4–10. Reproduced from: http://www.globalwarmingart. com/wiki/Image:Phanerozoic Carbon Dioxide png. (See companion website for color version.) provided empty ecological niches and, hence, evolutionary opportunities. The latter increased the rate of macro- and micronutrient transport into the soils, freshwaters, and ocean. One of the important innovations of the Cambrian biological explosion was the evolution of organisms capable of generating hard parts. Calcium minerals (carbonates and phosphates) rapidly emerged as a dominant building material used by shallow-water benthic invertebrates for shells and skeletons. During the Cambrian, various types of microbes, including cyanobacteria, contributed to the formation of extensive shallow-water calcified reef structures. The rise of large vascular land plants during the Devonian period accelerated the removalofCO2 from the atmosphere and the addition of O2. Land plants also acceler- ate chemical weathering rates, facilitating erosion of the continental land masses and increasing the rate of nutrient delivery to the oceans. The increasingly oxic atmosphere of the Devonian (Figure W8.4) stimulated further evolution of animals, leading to increased bioturbation of soil and sediments and lower burial rates of organic matter. Pelagic calcifying plankton rose to ecological prominence during the Mesozoic era following the largest of the mass extinction events of the Phanerozoic eon that occurred at the end of the Permian period (Figure W21.1). The environmental driving forces favoring the calcifiers are thought to have included: (1) high rates of nutrient supply due to enhanced continental weathering rates and (2) extreme oversaturation in the surface waters with respect to calcite. Dinoflagellates and diatoms also came to prominence at 8 Web Supplement 25.4

310 (Günz-Mendel) Yarmouth Sangamon Holocene Aftonian Aftonian (Mindel-Riss) Interglacial (Riss-Würm) Interglacial (I) (II) (I) (II) (III) Interglacial 290

270 Interglacial

250 (ppm) 2 230 CO

210 Glacial 190 “Independence” Nebraska Kansan(?!) IIIinois Wisconsin (Günz) (1) (2)(Mindel) (3) (4) (Riss) (Würm) 170 2 2 2 2 2 2 2 2 2 2 2 2 2 0 650000 600000 550000 500000 450000 400000 350000 300000 250000 200000 150000 100000 50000

Years before present

FIGURE W25.3

Atmospheric CO2 reconstructed from air bubbles in polar ice during the late Pleistocene. Note that

human activities have driven PCO2 levels to over 380 ppm, well above the scale on this graph. Ice core data sources: (1) (Combined) Law Dome: 1006 AD to 1978 AD, (2) Vostok: 2,342 to 417,160 years before present and (3) EPICA/Dome C: 415,000 to 650,000 years before present. For citations see: NOAA's World Data Center Ice Core Gateway at http://www.ncdc.noaa.gov/

paleo/icgate.html. Drawn by T. Ruen, http://en.wikipedia.org/wiki/Image:Atmospheric CO2 with glaciers cycles.gif.

this time (late Permian to early Triassic). Along with the coccolithophorids, these three groups of plankton are now responsible for most of the export flux of POC to the deep ocean and its sediments. The final element of the modern oceanic carbon cycle was established once biogenic calcite starting depositing onto the deep seafloor. This provided three important stabilizing feedbacks to the global carbon cycle. First, the presence of sedimentary CaCO3 stabilized deepwater carbonate concentrations, which provides buffering of atmospheric P on time scales of millennia. Second, accumulation of biogenic calcite CO2 increased the CaCO3 content of the sediments and, hence, enhanced the decarbonation component of the geological carbon cycle. Third, as described in Chapter W15.6, shifts in the CCD enable operation of a calcite compensation feedback that acts to stabilize atmospheric P by altering the spatial extent of marine sediments in which CaCO CO2 3 can be buried. Before the advent of the pelagic calcifiers, calcite buffering could occur only through adjustments in the degree to which shallow-water carbonate deposition took place. Because the shallow waters are supersaturated with respect to calcite, their sediments Web Supplement 25.4 9

always lie above the CCD. Hence, they cannot support a negative feedback loop that buffers against P change as effectively as can the deep-sea sediments. In the case of CO2 the latter, a decline in atmospheric P leads to a deepening of the CCD and, hence, CO2 an increase in the spatial extent of calcite deposition in the deep-sea sediments. As per Eq. 25.1, calcite deposition serves as a net source of CO2 to the atmosphere. Because the shallow-water sediments already lie above the CCD, a decline in P serves only CO2 to increase the degree of calcite supersaturation in their overlying waters. If the supersaturation becomes high, calcite precipitation rates increase, but this effect is likely countered by a concurrent reduction in the volume of the shallow waters and spatial extent of their underlying sediments. The latter is a consequence of lowered sea level associated with glaciation induced by global cooling during times of low P . CO2 To summarize, the biological innovations of the Phanerozoic have established a variety of feedbacks, some stabilizing and some destabilizing, that act on the atmospheric CO2 reservoir. Examples of how these feedbacks have operated during the major climate shifts of the Phanerozoic are described next following a short summary of how the global carbon cycle interacts with the atmospheric O2 and sulfur cycles.

Stabilization of the Atmospheric O2 by the Carbon and Sulfur Cycles

In Chapter W8.6, we discussed the ocean’s role in controlling atmospheric O2 levels. We now revisit and expand upon this topic by looking more closely at the linkages among the carbon, sulfur, and oxygen cycles. Sulfur is removed from the oceans via the formation of sedimentary pyrite and the precipitation of CaSO4. The burial of organosulfur compounds may also be important. Some of the sulfide fueling pyrite formation is hydrothermal in origin, but most comes from microbial sulfate reduction. The precipitation of CaSO4 occurs in hydrothermal vents and evaporites. The reduced sedimentary sulfur acts as a significant sink of O2 when tectonic uplift exposes pyrite to the atmosphere. The exposed pyrite reacts with atmospheric O2. This chemical weathering process oxidizes the pyrite’s sulfur and iron. Uplifted and exposed CaSO4 also undergoes chemical weathering in which congruent dissolution solubilizes dissolved sulfur as sulfate without affecting the atmospheric O2 reservoir. Any pyrite and CaSO4 that are not uplifted are carried back into the mantle via subduction. During this process, the sulfur is converted into gaseous form (H2S and SO2) and emitted back into the atmosphere via volcanic eruptions. Oxidation of this H2S and SO2 is another significant sink of atmospheric O2. Like the global carbon cycle, the sulfur cycle has undergone a massive reorganization due to the evolution of microbial life. In the case of sulfur, the two critical microbes are the sulfide oxidizers and the sulfate reducers. Recall that the latter form a microbial con- sortium with the anaerobic methane oxidizers, thereby providing an important linkage to the carbon cycle (Chapter 7.3.2). The paleoceanographic record indicates that both sulfate reducers and sulfide oxidizers evolved early in the Archean eon when the ocean was still anoxic. In an ocean without O2, sulfide oxidation would have taken place via anoxygenic photosynthesis in which sulfide (−II) is oxidized to elemental sulfur(0). 10 Web Supplement 25.4

The ocean provides an additional linkage among the sulfur, carbon, and O2 cycles. As discussed in Chapter W8.6, the burial of organic carbon is a net source of O2 to the atmosphere and its weathering (via aerobic respiration) is a net sink. Negative feedbacks involving sedimentary organic matter and pyrite serve to stabilize atmospheric P .For O2 example, increased P promotes an enhanced oxidation rate of uplifted sedimentary O2 organic carbon and pyrite, thereby consuming excess O2. This feedback operates only if the atmospheric and seawater levels of P are rate limiting. In the case of the former, O2 this requires that tectonic uplift and physical weathering be fast enough to keep up with rising P . In seawater, evidence for the controlling effect of O is provided by O2 2 the inverse correlation of O2 exposure time (OET) with sedimentary organic content (Chapter 23.7.2). Thus, the rate of loss of carbon from the oceans via burial in the sediments is dependent, at least in part, on O2 levels, which are in turn greatly inﬂuenced by the degree to which sulfur is partitioned among its reduced and oxidized reservoirs. 2+ The level of oxidation of seawater also inﬂuences CaCO3 precipitation rates as Mn and Fe2+ inhibit the formation of calcite. This provides yet another linkage between the carbon, sulfur, and O2 cycles. 25.4.4 Carbon Cycle Perturbations in the Phanerozoic

Some very large swings in P , P , and temperature have occurred during the Phanero- CO2 O2 zoic. Considerable attention is focused on understanding the causes of these swings and how the global carbon cycle responded in the hopes of obtaining insight into the likely consequences of our anthropogenic additions to the atmospheric CO2 reservoir. Our additions have driven atmospheric P from its preindustrial level of 280 ppm to CO2 386 ppm (circa 2009).2 The last time P was this high was 34 to 40 mybp. Since this CO2 last period of high P , Earth’s climate has shifted from an ice-free mode into one CO2 characterized by rapid cycling between glacial and interglacial conditions accompanied by oscillations in global temperature and atmospheric P . Some scientists hypothesize CO2 that our anthropogenic CO emissions will drive atmospheric P over a threshold that 2 CO2 will switch Earth’s climate out of these “rapid” glacial-interglacial cycles and into some other mode that will likely start with a large-scale deglaciation. While P may not in CO2 and of itself initiate global climate change, its remarkable covariance with temperature and ice volume suggests that it plays an important role in amplifying climate change. Events during the Phanerozoic support this hypothesis.

Mass Extinctions, Flood Basalts, OAEs, and the Carbon Cycle At present, paleoceanographers have identified four periods of significant climate change in the Phanerozoic. All were accompanied by major extinction events and, hence, have been designated by geologists to represent boundaries that distinguish geologic periods. The first of these climate change events occurred during the late Carboniferous and

2 This is the global monthly mean value for January 2009, after correction for seasonal effects, as reported by Dr. Pieter Tans, NOAA/ESRL (www.esrl.noaa.gov/gmd/ccgg/trends). Web Supplement 25.4 11

early Permian, 300 to 250 mybp. During the late Carboniferous, the longest and most severe glaciation of the entire Phanerozoic eon ended abruptly at 290 mybp. Within 10 million years, P rose from present-day levels to 3500 ppm (Figure W25.2)! The amount CO2 of carbon injected into the atmosphere was equivalent to our entire modern-day fossil fuel reservoir. This was accompanied by a steep drop in P , with levels decreasing O2 from an all-time high of 25% v/v down to 16% v/v (Figure W8.4). The high P led CO2 to global warming and deglaciation. Atmospheric P then underwent several short, CO2 high-amplitude oscillations in which intervals of low P supported transient periods CO2 of glaciation. This oscillating finally ended around 270 mybp with P levels stabilized CO2 somewhere between 2500 and 3500 ppm, leaving Earth’s climate at a relatively warm state that persisted until the end of the Mesozoic Era, 100 mybp. The beginning parts of this scenario are ominously similar to the current day, suggesting that we will eventually see a period of great climate instability followed by permanent “hot house” conditions. The next two significant bouts of climate change occurred across the Permian- Triassic boundary and the Triassic-Jurassic boundary. Both were accompanied by mass extinction events. The end Permian extinction is the largest on record. The Triassic- Jurassic extinctions were similar in scale to those of the Cretaceous, tying for second behind the end Permian extinctions. All three events coincide with periods of intense flood basalt formation (Figure W25.4). Some of these massive piles of frozen lava are millions of cubic kilometers in volume and, hence, are also called large igneous provinces (LIPs). The closest modern-day example of this type of eruption is the mantle hot spot that is building the Hawaiian Island chain (Chapter 19.2.3). Geologists hypothesize that Earth’s crust is built by two cycles: one driven by plate tectonics, i.e., the Supercontinent and Wilson cycles (Chapter 19.5.2), and one involving massive hot-spot- type eruptions that produces LIPs. The former generates oceanic crust and the latter contributes significantly to building of continental land masses. The largest of these flood basalts, the Siberian traps, were laid down 250 mybp at the Permian-Triassic-boundary. The Central Atlantic Magmatic province flood basalts, located in Central America, were deposited at the Triassic-Jurassic boundary and the Deccan traps, located in India, at the end of the Cretaceous. Exactly how these flood basalts might have caused massive die-offs is not known but is thought to result from a combination of factors such as (1) emissions of volcanic CO2 and other acid volatiles resulting in acid rain and (2) ejection of huge amounts of ash into the atmosphere. The latter would have reflected insolation and thereby stopped or reduced photosynthesis. The flood basalts were formed at times when large increases in P occurred. In the CO2 case of the Triassic-Jurassic boundary, some of the proxies used to reconstruct paleo-CO2 suggest levels as high as 6000 ppm! Such high CO2 levels led to a biocalcification crisis at the end of the Triassic as acidification of the ocean was intense enough to cause the collapse of carbonate reef platforms on the continental shelves. Global warming during these events was probably enhanced when rising temperatures eventually led to the melting of gas hydrates in the marine sediments, causing the release of methane. The last large warming event in the Phanerozoic occurred at the Paleocene-Eocene boundary and is called the Paleocene Eocene Thermal Maximum (PETM). The climate during the early Eocene (52 to 55 mybp) was the warmest of the past 65 million years. 12 Web Supplement 25.4

end Frasnian ? 350

300 Viluy (Siberia) Emelshan

Siberia end Guadalupian 250 end Permian Central Atlantic Magmatic Province

karoo and Farrar karoo end Triassic

200 Parana and Etendeka end Pliensbachian Rajmahal, Kerguelen, Ontong Java (phase 1) 150 end Jurassic?

Madagascar and Caribbean Plateau, Ontong Java (phase 2) end Valanginian?

Deccan end Early Aptian?

100

and Geological Time Scale Boundaries (Ma) end Cenomanian? Ethiopia and Yemen

Ages of Mass Extinctions, Oceanic Anoxia Events end Turonian? North Atlantic Volcanic Province (phase 1) Province North Atlantic Volcanic end Cretaceous Columbia 50 end Paleocene end early Oligocene?

end early Miocene?

0 50 100 150 200 250 300 350 Ages of Continental Flood Basalts or Oceanic Plateaus (Ma)

FIGURE W25.4 Age correlations between LIP, OAE, and major extinction events. The four largest recent mass extinctions and corresponding LIPs are shown by the large dots. Ma = million years. Source:From Courtillot, V. E. and P. R. Renne (2003). On the ages of food basalt events. C. R. Geoscience, 335, 113–140.

◦ This warming commenced 55.8 mybp, when global temperatures increased by 5 to 10 C over a period of 10 to 20 ky. Atmospheric CO2 rose as high as 1000 to 1500 ppm. This climate change caused rapid (10,000 y) changes in the species composition of the terrestrial flora. The PETM is thought to have been triggered by increased seafloor spreading along the Mid-Atlantic Ridge near Iceland. In addition to emitting volcanic CO2, this enhanced tectonic activity created a land bridge that altered deepwater circulation pat- terns by restricting the flow of water from the Nordic Seas to the North Atlantic. This led to warming of the deep waters and possibly the release of methane via melting of the gas hydrates buried in the marine sediments. The PETM is viewed as an important analog to present-day global warming as the rates and magnitudes of carbon release are similar to the present-day anthropogenic inputs. Web Supplement 25.4 13

3000 Paleosols 2500 Stomata Phytoplankton Boron 2000

1500 (ppm) 2

CO 1000

500

0 Cret. Paleogene Neogene 80 70 60 50 40 30 20 10 0 Time (Ma)

FIGURE W25.5

Atmospheric CO2 and temperature records for the late Cretaceous to present day (0 to 80 Ma). Cold periods with strong evidence for geographically widespread ice are marked with dark shaded bands. Cool-to-cold periods with indirect or equivocal evidence for ice are marked with light shaded bands. The horizontal lines at 1000 and 500 ppm CO2 represent the proposed CO2 thresholds for, respectively, the initiation of globally cool events and full glacials. Paleoceanographic

CO2 proxies are defined in the graph legend. Source: From Royer, D. L. (2006). CO2-forced climate thresholds during the Phanerozoic. Geochimica et Cosmochimica Acta 70, 5665–5675.

At the Eocene-Oligocene boundary (33.8 mybp), Earth’s climate underwent its latest large-scale shift, transitioning into a relatively cold period that has persisted to date. This shift is thought to be driven by a decline in seaﬂoor spreading rates and a concurrent increase in weathering rates, both of which led to lowered CO2 (Figure W25.5). The cooler temperatures caused a marked increase in ice volume in the earliest Oligocene, mostly in Antarctica. (The Arctic Ocean did not acquire its sea ice cover until the late Miocene, 10 mybp.) The formation of the Antarctic ice sheet is thought to have increased the rate of meridional circulation. Fast surface-water circulation rates and enhanced nutrient availability enabled the diatoms to come to ecological dominance. Because diatoms have a high export efﬁciency for POC, their rise caused the soft tissue pump to speed up, causing a further drawdown of CO2 and more cooling.

Fluctuations in Solar Radiation At some point during the Oligocene, P levels declined to levels low enough CO2 that changes in insolation could signiﬁcantly affect climate. This initiated the glacial- interglacial oscillations that have continued through to the present day with increase in amplitude in the Pleistocene (0 to 1.8 mybp). The insolation changes driving these glacial-interglacial cycles are caused by shifts in Earth’s distance from the Sun. These shifts are driven by periodic variations in four orbital parameters: (1) the elliptical shape of Earth’s orbit, (2) the tilt or obliquity of Earth’s axis relative to the plane of its orbit, 14 Web Supplement 25.4

(3) the wobble, or precession, of Earth’s axis, and (4) shifts in the inclination of the plane of Earth’s orbit relative to the rest of the solar system. The periods of these orbital variations are on the order of 100,000, 41,000, 23,000, and 70,000 y, respectively, and are called Milankovitch cycles. At any given time, the amount of insolation received by the Earth can be predicted by considering the combined effects of the four orbital parameters. Since the Oligocene, the Milankovitch cycle with the 100,000-y period has demonstrated the strongest correlation with global climate change, i.e., the timing of glacial and interglacial conditions. Other potential drivers of insolation changes that are large enough to affect climate include changes in atmospheric albedo and in solar irradiance. Various natural phenomena increase atmospheric albedo. These include volcanic emissions of gas and ash. For example, the volcanic gas, sulfur dioxide, is converted into sulfuric acid aerosols that reflect radiation. The cooling effect of volcanic eruptions is shown in Figure 25.8a. Other important natural sources of reflective particles are: (1) aeolian dust transported from deserts and semiarid regions and (2) smoke from forest fires. The intensity of these sources is linked to climate. Changes in solar irradiance are thought to occur over many time scales with the most well known being the sunspot cycle. During historical times, sunspot activity has been observed to oscillate with an average period of 11 years. During the low part of the cycle, little-to-zero sunspot activity occurs. Interruptions to this cycle have involved extended periods of little-to-no sunspot activity, i.e., the Wolf (1280–1350 AD), Sporer¨ (1450–1550 AD), and Maunder (1645–1715 AD) minima. These minima coincided with a period of low atmospheric temperature called the “Little Ice Age” (see Figure W25.6). Conversely, an extended period of high sunspot activity, called the Medieval Maximum

Medieval warm period 2 Heinrich event 1 C) 8 Holocene maximum

158C 0

Little Ice age 22 Younger Dryas

Change in temperature ( Change in temperature 24 Pleistocene epoch Holocene epoch

18 16 14 12 10 8 6 4 20 Thousands of years before present (B.P.)

FIGURE W25.6 Temperature history of Earth during the last 18,000 y. See Chapter 13.5.4 for a description of Heinrich events. Source: From Mackenzie, F. T. (1998). Our Changing Planet: An Introduction to Earth System Science and Global Environmental Change, 2nd ed. Prentice Hall (Fig. 11.6a on p. 162). Web Supplement 25.4 15

(1100 to 1250 AD), coincided (at least partially) with a period of warmth called the Medieval Warm Period. In more recent times, extended periods of lower sunspot activity have taken place, namely from 1885 to 1910 and 1960 to 1975. Although these drops have not been as pronounced as the minima that occurred during the Little Ice Age, they have coincided with periods of modest cooling. Theoretical calculations indicate that the declines in sunspot activity have not been, in and of themselves, sufﬁcient to cause the observed cooling effects. Scientists hypothesize that positive feedbacks could amplify the effects of reduced solar irradiance on global climate, thereby providing a potential causal linkage between solar irradiance and global climate. One such proposed feedback involves a reduced photochemical production of ozone in the stratosphere when solar irradiance is low. Since ozone is a greenhouse gas (GHG), its lowered production enhances the cooling effect of reduced solar irradiance. Evidence exists for longer time-scale variations in solar irradiance. As noted earlier, the 11-y sunspot cycle has been periodically interrupted on centennial time scales. Since the cause of this interruption is not known, prediction of the timing of the next Maunder-type minimum is not possible. On millennial time scales, evidence for changes in solar irradiance have been obtained from the polar ice record. Downcore variations in the 10Be content of the ice are interpreted as a record of changes in solar irradiance because 10Be is produced in the atmosphere by chemical reactions whose production rates are affected by solar irradiance. The ice core 10Be record goes back 9500 y and indicates that solar activity has been highly variable throughout the Holocene. Two maxima in solar activity have occurred, one at 8000 ybp and the other 2000 ybp. Since that time, solar activity has generally been declining.

Ice Ages in the Pleistocene Over the past 1.8 million years, Earth has cycled in and out of Ice Ages during which large continental ice sheets have waxed and waned. These changes in ice volume and regional temperatures are well correlated with oscillations in the greenhouse gases, CO2, CH4, and N2O (Figure W25.7). As in earlier times, high GHG levels have coincided with warm and interglacial conditions. Conversely, cold and glacial conditions have coincided with low GHG levels. During the late Quaternary, Earth has been cycling in and out of Ice Ages every 100,000 years or so, corresponding to the Milankovitch ellipticity orbital cycle (see Figure W25.3). Lower amplitude swings in ice volume have been occurring at a fre- quency of 41,000 y, matching the period of the tilt variations. Over the past 650,000 y, Earth has experienced seven Ice Ages and seven interglacials, including the one that we are in now. While the ice core data show us that some periods of rapid temperature change were not accompanied by GHG changes, all periods of rapid and signiﬁcant atmospheric P change have been accompanied by temperature swings. Similar behav- CO2 ior is seen in the other GHGs, CH4 and N2O. In the case of methane, atmospheric concentrations are not as well correlated with climate shifts. The role of the GHGs in climate change is still a matter of hot debate and research effort. While the Milankovitch cycles are the major agent of long-term climate change, 16 Web Supplement 25.4

1600

1200

Methane 800 Methane (ppb) 400 350 Carbon Dioxide 300

250

200 300 Carbon dixoide (ppm)

280 Nitrous Oxide 260 240 220 Nitrous oxide (ppb) Nitrous 200 Marine Isotope Stage 2380 11 12 13 14 15 16 dD 2400

2420 D ‰) d ( 2440 Temperature proxy Temperature 0 100 200 300 400 500 600 Age (years) 3 103

FIGURE W25.7

The greenhouse gas (CO2,CH4, and N2O) and deuterium (␦ D) records for the past 650,000 years from EPICA Dome C and other ice cores, with marine isotope stage correlations (labeled at lower right) for stages 11 to 16. ␦ D, a proxy for air temperature, is the deuterium/hydrogen ratio of the ice, expressed as a per mil deviation from the value of an isotope standard. More positive values indicate warmer conditions. Data for the past 200 years from other ice core records and direct atmospheric measurements at the South Pole are also included. Source: From Brook, E. J. (2005). Tiny bubbles tell all. Science 310, 1285–1287.

the GHGs undoubtedly play an important role in determining how Earth’s climate responds to changes in insolation. One of the key features of the atmospheric P , CO2 P , and temperature oscillations is their slow decline into glacial conditions followed CH4 by a rapid return to high levels during the interglacials. This suggests that a threshold is reached that triggers a rapid shift from glacial to interglacial conditions. It also suggests that whatever drives Earth into an Ice Age is not necessarily the reverse of what happens to end the Ice Age and bring Earth back to interglacial conditions. Web Supplement 25.4 17

Most scientists agree that oscillations in the GHG concentrations over Earth’s history have affected past climates through a variety of physical, chemical, biological, and geological feedbacks, some positive and some negative. Various triggers and feedbacks that are thought to be involved in the Pleistocene’s glacial-interglacial oscillations are sum- marized in Table W25.1. Many have been discussed earlier as they involve oceanic processes. For example, high atmospheric dust loadings are a common characteristic of maximum glacial conditions (Figure 5.13). This suggests an important role of dust in (1) reflecting insolation and (2) enhancing the drawdown of atmospheric P . The CO2 latter is achieved through two effects: (1) the dust increases the delivery of trace metals, thereby by relieving the micronutrient limitation of marine plankton (Figure 11.9) and (2) the dust serves as ballast to the sinking flux of POC, thereby increasing the export efficiency of POC. Although the glacial-interglacial cycling in the late Quaternary is well correlated with oscillations in temperature and the GHGs, these oscillations have been modest compared with those of earlier times. Atmospheric CO levels have varied by only 80 2 ◦ to 100 ppmv and methane by 350 ppb, while temperature has shifted by only 10 C between the glacial and interglacial states. This suggests that the Earth’s carbon cycle and climate have been stabilized by powerful negative feedbacks that prevent the occurrence of runaway ice houses or runaway hot houses as seen in earlier times. The most important of these feedbacks are thought to involve the GHGs and ice area. It is notable that relatively small shifts in GHGs and ice cover seem to result in enough of a climate change to shift Earth between glacial and interglacial states. This suggests that the human-induced changes in atmospheric P could have a substantial CO2 impact on the timing of our planet’s next Ice Age. Based on the past 100,000-y periodicity of the glacial-interglacial cycling in the late Quaternary, Earth should be entering or close to (within a few thousand years or so) another period of glaciation. But all bets on this are now off, given that anthropogenic forcing has driven P to levels last seen CO2 30 million years ago during the PETM. As discussed in Chapter 25.5, humans are now the most likely drivers of global climate change. No single process is likely responsible for causing Earth to shift between glacial and interglacial states. Current hypotheses include a series of progressive steps over which various positive feedbacks successively kick in. An example is given in Figure W25.8 in which intermediate conditions are envisioned as preceding full-blown glacial or interglacial states. The progression starts with an insolation change associated with one or more of the Milankovitch cycles. The importance of this change is not in its overall impact on global heating or cooling, but rather on an alteration in seasonality. Specif- ically, a descent into an Ice Age is promoted by cooler summers that prevent melting of ice that accumulated during winter. Changes in ice cover can lead to changes in nutrient delivery to the ocean via terrestrial runoff and ocean circulation. These in turn affect the marine biota. The combined physical and biotic changes will, in turn, alter the rate at which the ocean’s carbon pumps can operate. The scenario illustrated in Figure W25.8 starts with a physically driven descent into glacial conditions. The ocean enables an initial P drawdown through two CO2 18 Web Supplement 25.4 -rich deep waters). 2 Regions with underutilized nutrientries invento- (primarily the Southernorial Ocean, Pacific equat- and NorthSouthern Pacific, Ocean with containing the aunused majority nutrients of having awith strong CO connection Mainly in regions dominatedophorid by production coccolith- today butshift a expected. global Nutrient-limited regions (e.g., low-latitude gyres). Regions north of thethat Antarctic are Polar silicon Front limited.enced Low by latitudes mixing influ- ofwater subantarctic into mode the thermocline. Regions influenced by aeolianand dust with input underutilized nutrientinventories (N, (Southern P, Ocean, Si) Northequatorial Pacific, Pacific). . 2 2 from con- 2 relative to organic 3 . Increased nutrient utiliza- 2 to surface waters. 2 . 2 marine productivity, carbon uptakewaters, in and subsequent surface carbondeep flux ocean, to which the lowers atmospheric CO nutrient limitation andin allows marine biomass global anddeep increase carbon ocean. export to in surface waters removes CO change. Iron-fertilized diatoms inOcean the Southern take up lessto silicic nitrate. acid The relative unused silicic acid is exported carbon from surface totains high deep surface-water ocean alkalinity main- (making CO tact with the atmosphereatmospheric and CO results in lower more soluble) and reducingcarbonate deep-ocean burial. Associated whole-ocean alkalinity increases, reducing atmospheric CO tion could occurproduction either or by by increasing decreasingmixing surface vertical and therefore theand supply CO of nutrients Oceanic Processes and Their Postulated Roles as Triggers and Feedbacks in Global Climate Change. Table W25.1 Postulated FeedbackMechanisms BIOTIC PHENOMENA Iron fertilization DescriptionWhole-ocean nutrient increase Increased ocean nutrient Increased reservoir inputs alleviates of iron-richNutrient dust utilization increase Oceanic Regions Influenced Increased utilization of carbon andPIC-to-POC nutrients rain ratio Decreased export of CaCO Silica leakage Mechanism by which the rain ratio could Web Supplement 25.4 19 308, Science cycles. Global Biogeochemical 2 Subpolar and polar oceans. Coastal waters. Whole ocean. Whole ocean. 2 can 2 retention. 2 -to-organic carbon rain ratio. . Such changes would eventually 3 2 It also alters nutrient transport. temperature changes. Changes inalter salinity. ice The volume effect of temperature on CO rich shelf sedimentsincreases to delivery of terrestrial nutrientsto weathering and the alkalinity ocean. Asdelivery sea declines. level The increases, exportPIC this of reburies POM nutrients and andshelf alkalinity sediments. into the which aaltered change terrestrial in input, affectsof alkalinity, the CO such solubility as from and sea ice coverage.waters Mixing control rates the in degree polar to which CO to silicon-limited regions, increasing diatom production at thecoccolithophorids, expense and of decreasing the CaCO solubility is opposite tocountering that some of of salinity, thereby the temperature effect. alter preservation of carbonatevia sediments carbonate compensation asChapter per W15.6.1. be retained. Low mixingdensity due stratification to increases enhanced CO After Kohfield, K. E., C. Le Quere, S. P. Harrison, and R. F. Anderson. (2005). Role of marine biology in glacial-interglacial CO 20, GB2010. : PHYSICAL PHENOMENA Temperature and salinity changes Milankovitch cycles alter insolation leading to Shelf sedimentsChange in ocean As alkalinity sea level declines, exposure of organic- Considers just the chemical response in Changes in polar circulation Caused by changes in temperature,Source salinity, 74–78; and Peacock, S.,Cycles E. Lane, and J. M. Restreo (2006). A possible sequence of events for the generalized glacial-interglacial cycle. 20 Web Supplement 25.4

280

270 Reduction in temperature Increase in low-latitude and northern temperature Reduction in high-latitudemixing 260 Increase in northern hemisphere mixing

Decrease in mean-ocean alkalinity and phosphate 250 Rise in sea-level; decrease in mean-ocean salinity 2

CO 240 P

230

220 Atmospheric Increase in Southern Ocean temperature Increase in mean-ocean alkalinity and phosphate 210 Increase in southern Ocean mixing Fall in sea-level; rise in mean-ocean salinity 200

190 Interglacial Glacial Intermediate Interglacial 180 20 40 60 80 100 120 Time (thousands of years)

FIGURE W25.8

Atmospheric PCO2 over the last glacial-interglacial cycle. Included are short descriptions of the mechanisms thought to have driven climate change in each step of the cycle. Source:From Peacock, S., E. Lane, and J. M. Restreo (2006). A possible sequence of events for the generalized glacial-interglacial cycle. Global Biogeochemical Cycles 20, GB2010.

processes: (1) enhanced GHG solubility due to cooling of surface waters and (2) reduced ocean mixing at high latitudes. The latter enhances CO2 retention in the deep waters, particularly in the Southern Ocean. Falling sea level exposes organic-rich shelf sediments to terrestrial weathering, thereby increasing nutrient and alkalinity delivery to the ocean. This stimulates the marine biota to produce more POC and PIC, both of which are then exported to the deep sea and sediments. Burial of POC and PIC further enhances the P drawdown and Earth’s descent into a full-blown glacial state. CO2 The return to interglacial conditions is envisioned as starting with a Milankovitch-driven increase in temperature that increases mixing in the Southern Ocean, thereby reducing its ability to retain CO2. As sea-level rises and inundates the continental shelf regions, the delivery of nutrients and alkalinity from terrestrial weathering diminishes. This and warming act to reduce the rate at which the oceanic carbon pumps operate, causing a continuing rise in atmospheric P . This provides a positive feedback to global CO2 warming that continues until full-blown interglacial conditions are attained. Likely important details in these scenarios include shifts in plankton community structure induced by changes in nutrient and micronutrient availability. Such shifts Web Supplement 25.4 21

influence the efficiency and, hence, the magnitude of the biological pumps. Particu- lar attention has been focused on iron availability as this provides a linkage to climate and ocean circulation. Climate controls the aeolian transport of iron. Ocean circulation controls the lateral transport of sedimentary iron resuspended from shelf sediments. Transport of this iron into the euphotic zone requires deep mixing by winter storms. In both cases, solubilization of the particulate iron is required to make this micronutrient bioavailable. The speciation calculations presented in Chapter 5.7 are illustrative of efforts now underway to predict how much bioavailable iron can be generated from the particulate forms. Because iron reduces the silica requirement of diatoms, iron fertilization of the Southern Ocean could be a key agent of control for the biological carbon pumps (Chapter 16.6.2). Plankton are also important climate control agents as they are the source of gases that are atmospherically active, such as DMS and the halocarbons (Figure W23.1 and Chapter 22.4.10). Many of the ocean processes that are likely to exert global impacts on climate either drive, or are driven by, meridional overturning circulation. For example, changes in the rates of deepwater formation and its return to the sea surface should affect the rate at which nutrients are returned to the euphotic zone, hence altering the export flux of PIC and POC. For example, a major slowdown in meridional overturning circulation is posited to have occurred 11,500 to 12,900 ybp as a result of a significant cooling ◦ event (about 2 C) called the Younger Dryas (Figure W25.6). This cooling was a brief interruption of the general warming that has been ongoing since the last Ice Age ended 18,000 ybp (Wurm¨ glaciation). The cooling during the Younger Dryas is thought to have been caused by a disrup- tion in NADW production caused by the release of glacial meltwaters into the North Atlantic. As Earth warmed, glacial meltwaters pooled in large lakes located at the present site of Northern Canada. Once the ice retreated sufficiently northward, the meltwaters were free to flow into the North Atlantic. The flow path seems to have been across the Great Lakes, down the St. Lawrence River, and into the Labrador Sea, one of the present-day sites of NADW formation. The presence of large amounts of freshwater should have reduced surface-water salinities. Since the temperature at which seawater can attain its highest density is inversely related to its salinity, fresher waters cannot get as dense as saltier waters (Figure 2.12c). Thus, despite continued winter cooling, deep water formation would have ceased and led to regional atmospheric cooling for two reasons. First the heat that would have otherwise been released to the atmosphere by the sinking of deep water was instead retained by the low-salinity surface waters. Second, the surface pool of freshwater prevented the warm waters of the Gulf Stream from reaching subpolar latitudes. Reestablishment of meridional overturning circulation took place 1000 y later, probably as a result of cooling that enabled regrowth of the ice sheets, thereby halting the flow of meltwater to the North Atlantic. The cause of this cooling is unknown. The formation of polar ice in the Northern Atlantic is also considered an important control of meridional overturning circulation, making it another likely driver of rapid climate change in the late Quaternary. This linkage is of particular concern as the rate of loss of polar ice in the northern hemisphere is accelerating. 22 Web Supplement 25.4

The power of meridional overturning circulation in altering global climate derives from it ability to control heat transport through the world’s oceans and, thus, broad- cast energy impacts worldwide. Other oceanic processes are thought to have a similar potential. They include: (1) the lateral spread of meltwaters on the sea surface in the form of eddies because of their potential to disrupt how the geostrophic currents transport heat and (2) changes in tropical rainfall rates as this affects the degree to which surface salty water can be returned from the surface waters of the Pacific and Indian Oceans back to the site of NADW formation in the North Atlantic. Warming in the Holocene The cooling event of the Younger Dryas (10,000 ybp) was the last gasp of the Wurm¨ glaciation. The climate changes that have ensued during the continued transition into an interglacial condition are illustrated in Figure W25.6. The step shift into warmer temperatures around 10,000 ybp defines end of the Pleistocene and the beginning of the Holocene epoch. As in past transitions from a glacial to an interglacial state, warming has proceeded at varying rates. The major features have been a period of rapid warming from 9000 to 5000 ybp called the Holocene Climatic Optimum during which ◦ temperatures were 0.5 to 2 C warmer than in the present day. This was followed by a period of minor glaciation from 5000 to 2000 ybp that was in turn followed by two short and small warming events. The one that lasted from 900 to 1300 AD is called the Medieval warm period. A similarly short and small cooling event, called the Little Ice Age, occurred from 1300 to 1850 AD. Since that time, global temperatures have been rising (Figure 25.8). During these low-amplitude temperature shifts, atmospheric GHG levels have been increasing, with higher rates of rise since the 1850s (Figure W25.9). Human activities could have contributed to at least some of the CO2 rise that occurred prior to 1850. This impact is postulated to be a possible result of land-use change, including deforestation and plowing, and biomass burning (forest fires). It is more likely that the pre-1850 rise in CO2 was primarily due to “natural” processes in which the ocean’s calcite compensation played an important role as per the following scenario. Warming at end of the last Ice Age melted continental ice, enabling regrowth of forests and wetlands. Most of the carbon fueling this regrowth was provided by the degassing of oceanic CO2 in response to a dampened solubility pump. This degassing led to a rise in carbonate ion concentrations in seawater by driving the following reaction toward the products: − → 2− 2HCO3 CO3 +CO2 +H2O (W25.1) High carbonate ion concentrations moved the CCD to deeper depths, leading to an increased rate of CaCO3 burial. This enabled release of even more CO2 from the ocean to the atmosphere. Once the terrestrial biomass stabilized and atmospheric drawdown of CO2 ended, the ocean was no longer driven to supply CO2 to the atmosphere and the CCD returned to shallower depths. Another interesting feature of the increase in P over the Holocene is that the rise CO2 has been sustained over a longer period than seen in the past three interglacials. As shown in Figure W25.3, these interglacials have been characterized by an initial short Web Supplement 25.4 23

Time (before 2005) Time (before 2005) (a) 10000 5000 0 (b) 10000 5000 0 2000 400 2000 ) ) 2 2 350 2 1500 2 0.4 350 1500 1000 300 1 500 1800 1900 2000 1800 1900 2000 Year Year 0.2 1000 300 Methane (ppb) 0 0 Carbon dixoide (ppm) Radiative forcing (Radiative forcing W m Radiative forcing (Radiative forcing W m 500 250

100005000 0 10000 5000 0 Time (before 2005) Time (before 2005)

300 2

300 270 0.1

240 1800 1900 2000 Year

270 0 Nitrous oxide (ppb) Nitrous Radiative forcing (Radiative forcing W m

10000 5000 0 Time (before 2005)

FIGURE W25.9 Atmospheric concentrations of (a) carbon dioxide, (b) methane, and (c) nitrous oxide over the past 10,000 years (large panels) and since 1750 (inset panels). The corresponding radiative forcings are shown on the right-hand axes of the large panels. Source: From IPCC Working Group I (2007). Climate Change 2007: The Physical Science Basis, Contribution of Working Group I to the Fourth Assessment Report of the IPCC. Cambridge University Press.

period of high CO2. The last interglacial with a similarly long period of sustained high CO2 occurred 400,000 ybp. A possible explanation for this lies in the 400,000-y periodicity of the combined effects of some of the Milankovitch cycles. On this time scale, some of the orbital parameters are out of phase. This results in a combined effect on insolation that is smaller than usually seen during the typical 100,000-y glacial-interglacial cycle. Figure W25.3 also shows that first two interglacials in the late Pleistocene did not have a well-defined initial increase in CO2 in contrast to the past five interglacials. This switch in system response is attributed to changes in oceanic circulation and sea-surface temperatures in the Southern Ocean. The last Ice Age terminated 10,000 to 15,000 ybp and has been followed by a period of general warming termed the Holocene Epoch. Prior to 1700, atmospheric CO2 levels during the Holocene varied by no more than 10 ppmv, fluctuating between 275 ppm and 285 ppm. The increase in P that has taken place from 1700 to 2000 has been CO2 24 Web Supplement 25.4

about 85 ppm, equivalent to 175 Pg C, or 30% of the preindustrial level. Atmospheric CO2 concentrations reconstructed from ice cores extending back 650,000 y have established that in the Late Quaternary, the typical glacial-interglacial difference has been ◦ 100 ppm, and was accompanied by temperature swings of about 10 C (Figure W25.3). Thus the present-day P level, which now exceeds 380 ppm, represents a large depar- CO2 ture from that of the past 650,000 y (Figures W25.3 and W25.9). Most of this increase is attributed to human activities as our known emissions are more than adequate to explain the observed atmospheric increase (Figure 25.5). As shown in Figures 25.4 and W25.9, the time frame over which atmospheric P has risen rapidly coincides with CO2 the period over which anthropogenic emission rates have increased. Figure 25.4b also documents that the changes in P have been largest in the northern hemisphere, CO2 matching the geographic source of our emissions. During this same period, the 14C content of the atmospheric CO2 has decreased, further supporting a fossil-fuel source. Figure W25.9 documents that the concentrations of other GHGs have also increased over the past century. Methane levels have roughly doubled. Although still in the ppb range, methane’s higher GWP makes its climate-forcing effect equivalent to 30% of that currently contributed by CO2. Fortunately methane is prone to oxidation, making its atmospheric lifetime relatively short (8 to 10 y). The increase in its concentration over time follows a temporal trend similar to that of CO2 until 1985. After 1985, the annual growth rate of methane has declined such that atmospheric levels have now stabilized.3 More than 50% of present-day global methane emissions are anthropogenic in origin, coming from fossil-fuel production (escape of natural gas from leaky pipelines), livestock (cattle), rice cultivation, and waste handling (including animal waste, domestic sewage, and landﬁlls). Biomass burning has also been an important source, particularly from 0 to 1000 AD. Natural emissions of methane are primarily from wetlands, peatlands, and tropical rain forests and from natural gas seeps associated with fossil-fuel deposits and the melt of gas hydrates. Swings in atmospheric methane levels during the end of the last glaciation are thought to be due to changes in wetland production rather than methane-hydrate destabilization.

3 This stabilization is thought to have been caused by the drying of wetlands, leading to decreased natural emissions and thereby countering some of the anthropogenic input. In 2007, P rose for the CH4 ﬁrst time since 1998, jumping from an annual global mean of 1775 to 1783 ppb. Likely sources are increased anthropogenic emissions from rapidly industrializing Asia and melting permafrost in the Arctic.