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Web Supplement 25.4

25.4 PAST VARIATIONS IN CLIMATE AND ATMOSPHERICCARBON DIOXIDE LEVELS

The geologic carbon cycle was described in Chapter 25.1 as providing the negativefeedback loop responsible for stabilizing PCO2

over time scales of 105 to 107 y. Inthis cycle, the burial of CaCO3 and SiO2 in ocean sediments fuels decarbonation reac-tions that produce CO2 and CaSiO3. The CO2 is degassed through volcanoes and theCaSiO3 is uplifted onto land. Weathering of uplifted CaSiO3 by atmospheric CO2 suppliesbicarbonate, silicate, and calcium ions to river runoff. In the modern ocean, planktontransform these ions back into CaCO3 and SiO2. If the rate of weathering is limitedby PCO2

, increasing PCO2should lead to enhanced uptake rates as a result of enhanced

weathering, thereby providing a negative feedback that acts to stabilize PCO2and the

sizes of the other crustal carbon reservoirs. Nevertheless, large transient swings in PCO2

have occurred during various periods of Earth’s history. A few have led to long-termreorganizations of the global carbon cycle. Paleoceanographers are particularly inter-ested in studying these events as they provide some clues as to what changes we canexpect as a result of our anthropogenic perturbations.

A key to understanding these past swings is recognizing the important role oftectonism in the geologic carbon cycle. First, it determines the rates and loca-tions of the decarbonation reactions and of uplift. Second, it determines the spatialextent of shallow shelves. The latter was critically important prior during the Pre-cambrian and first half of the Phanerozoic because carbonate burial was restrictedto the shallow shelves. In the mid-Phanerozoic, the evolution of pelagic calcifiersenabled burial of carbonate in deep-sea sediments. As discussed in Chapter 15.6,deep-sea sedimentary carbonate participates in a stabilizing feedback loop called cal-cite 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 primar-ily in the Pacific Ocean. Thus, the stability afforded by the geologic carbon cyclewill occur only over time scales long enough to capture a supercontinent cycle. Thestabilizing 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 andsubduction.

Tectonism is also important in the geological carbon cycle because it determinesland-mass distributions. In the present day, most of the land mass is located in north-ern hemisphere mid-latitudes. A land-mass distribution in which most of the continentsare located at tropical latitudes would be expected to result in a cold climate due toenhanced continental weathering rates. Conversely, our present geography should sup-port a relatively warm climate. This was not the case, at least prior to human-inducedwarming. The probable explanation is that the geological carbon cycle’s stabilizing feed-back is not presently working well. As noted earlier, most subduction is occurringin the Pacific where the sediments have a low carbonate content, thus disabling thegeologic carbon cycle’s metamorphic decarbonation linkage. On the other hand, thepresence of large land masses in the mid-latitudes does provide some climate stabilityin that an extreme cooling event leads to the formation of continental ice sheets. Thisslows continental weathering, permitting volcanic CO2 to reaccumulate and warm theatmosphere.

In the following subsections, the long-term evolution of the global carbon cycle isdiscussed along with the most likely explanations for the large transient swings in PCO2

that have occurred sporadically throughout Earth’s history. The marine carbon cycle hasplayed an important role in the past and is likely to determine the degree to which ourperturbations 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 sincethe early Precambrian. As shown in Figure W25.1, PCO2

is estimated to have droppedby a factor of 100 to 10,000 over Earth’s history, indicating a large-scale relocation ofcarbon 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 shownin Eqs. 25.1 through 25.8 stabilized PCO2

through the formation of carbonates via abio-genic precipitation in warm, shallow-water environments. This stabilizing feedback wasimportant as the Sun’s solar luminosity has been increasing (slowly) over time. On theearly Earth, solar luminosity was only 70% of the present day, so a long-term increase ininsolation should lead to an overall decline in PCO2

assuming that warming temperaturesenhance weathering rates.

The advent of life around 3.8 bybp added a new crustal reservoir to the global carboncycle, 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 theoverall trend of declining PCO2

. One of the earliest life forms were the methanogenicarchaeans. These microbes are thought to have converted virtually all of the primordialvolcanic H2 in the atmospheric to methane via reaction with atmospheric CO2. Methaneis 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 icefree. Prior to the evolution of methanogens, PCO2

levels were not high enough to keep

Web Supplement 25.4 3

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Ocean-covered earth

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Atmospheric PCO2relative 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 addedan important element to the carbon cycle that acted on global climate.

Methane production by the methanogens did not lead to a runaway greenhouseeffect, as a negative feedback was established through the interaction of UV radiationwith CH4. At high PCH4

, UV radiation induces the formation of a photochemical hazethat reflects insolation. This methane thermostating ended around 3.0 bybp with theadvent of oxygenic photosynthesizers. Their production of O2 led to the oxidation ofatmospheric CH4. With the loss of this potent greenhouse gas, the planet entered aperiod (mid-to-late Proterozoic) in which several global glaciation events occurred.

These are referred to as “ice-houses” or “Snowball Earths.” As discussed later, severalother causative factors contributed to the occurrence of these ice-house conditions. Incontrast to the climate variability of the Proterozoic, no Snowball Earths have occurredsince the beginning of the Phanerozoic 550 mybp. The switch back to climate stabilityis attributed to the advent of multicellular lifeforms and, in particular, to the calcitecompensation 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, accompaniedby banded iron formations (oxidized iron) overlain by very thick carbonate layers,


called cap carbonates.1 These carbonate deposits are typically 3 to 30 m thick andoccur on platforms, shelves, and slopes worldwide. The global distribution of theseunique sediments suggest that during a Snowball Earth episode, all the land masses andmost, if not all, of the oceans were covered by ice. The first series of Snowball Earthsoccurred between 2.2 and 2.45 bybp at the beginning of the Proterozoic (Huronianand Makganene glaciations). The second series took place during the late Proterozoic(the Sturtian glaciation 710 mybp, the Varanger-Marinoan 635 mybp, and the Gaskiers582 mybp). This latter period is referred to as the Cryogenian. Evidence suggests thatthe entire oceans froze during the Cryogenian. Life probably survived under thin ice atequatorial latitudes and at hydrothermal vents.

Entering Snowball ConditionsVarious combinations of four factors are thought to have led to the runaway cooling thatproduced Snowball Earth conditions in the Proterozoic. Briefly, these factors are (1) highweathering rates, (2) increased volcanic activity leading to the formation of continentalflood 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. (Thelast would have served to cause only the first series of Snowball Earths.) Once coolinghad started, a positive feedback is postulated to have occurred when the ice sheetsreached some critical latitude due to their cumulative high albedo, i.e., the ice sheetsreflect insolation back into space, thereby preventing its absorption by the greenhousegases.

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 ofsupercontinents; and (3) the formation of supermountain chains. During the period ofthe Snowball Earths, little continental area was located at high latitudes. Instead largepolar sea-ice caps were present. They reflected solar radiation but did not cover muchland area, leaving the continents ice free and susceptible to weathering. Immediatelypreceding the Cryogenian, around 830 mybp, a supercontinent, called Rodinia, beganbreaking up. This continued for nearly 200 my. Weathering rates increase when super-continents break up. When a supercontinent exists, most land area is far from theocean and therefore very dry. Conversely, when a supercontinent breaks up into smallfragments, formerly arid regions become wetter and, hence, weathering rates increase.Around 650 mybp, just prior to the Minoan glaciation, land fragments generated by theRodinian breakup collided to form the supercontinent Gondwanaland and gave rise toa 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 tothe ocean and is thought to have played a role in the subsequent evolution of multicel-lular organisms and other eukaryotes. All of these types of enhanced weathering leadto a drawdown in PCO2

and, hence, contributed to global cooling.

1 The formation of banded ironstones is discussed in Chapter W8.6.1.


At least some of the Snowball Earths seem to have been preceded by massivecontinental eruptions of basalt lava, called flood basalts. This happened on land massesvery close to the equator and, hence, was subject to intense weathering. This typeof basaltic rock weathers rapidly, acting as a rich source of Ca2+ and bicarbonate tothe ocean. When PCO2

levels dropped at the end of each Snowball Earth episode,

precipitation of this Ca2+ and bicarbonate led to the formation of the massive capcarbonates.

Finally, cooling events are also thought to be driven by tropospheric dust loadingcaused by encounters with interstellar giant molecular clouds (GMCs). Our solar systemencounters a dense GMC (>2000 H atoms cm−3) every 109 years or so. Average densityGMCs are encountered every 108 years, with greatest probability every ∼140 my as thisis the periodicity in which the solar system crosses the galactic spiral arms where GMCsare concentrated. The reflection of insolation by this dust is sufficient to cause SnowballEarth conditions, especially as the dust loading occurs too rapidly to be compensatedfor by the geological carbon cycle stabilizing feedbacks. Based on this periodicity, Earthhas likely encountered approximately four high-density GMCs, each of which could havetriggered 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 SnowballOnce a Snowball Earth condition began, cooling would have lowered the silicate weath-ering rate, stabilizing the climate system at the new colder state. Two other factorswould have contributed to lowered weathering rates: (1) at low PCO2

, the acidity ofrainwater would be reduced, and (2) the coverage of land by ice prevents weather-ing. Further stability would be provided by the high albedo contributed by the globalice sheets. Nevertheless, these feedbacks would eventually be overcome by the Geo-logical Carbon Cycle once enough time had elapsed for the decarbonation of calciumsilicate and the oxidation of organic matter buried in the sediments prior to the onset ofthe snowball condition. Continuing volcanic activity would have eventually resuppliedCO2 to the atmosphere along with iron to the deep sea through hydrothermal emis-sions. Conditions of lowered sea level are thought to increase the iron-to-sulfur ratio inhydrothermal emissions. Because the ice isolated the ocean waters from the atmospherefor millions of years, the bottom waters likely became anoxic. This promoted elevatedFe2+ concentrations as a result of anoxic conditions and insufficient sulfide for removalvia pyrite precipitation.

Once volcanically supplied PCO2rose to some critical level, warming must have

ensued. Melting of the global ice cover is thought to have happened quickly due tovarious positive feedbacks. For example, a rapid reduction in albedo would result fromthe pooling of meltwater on ice. Climate modeling suggests that the meltdown couldhave occurred in as little as 2000 y, leading to a variety of cataclysmic effects. First wouldhave been a rapid and large increase in the rate of river runoff. The flow of freshwaterinto the ocean would have led to an initial density stratification in which low-densityoxic meltwater formed a cap over the denser anoxic seawater. The meltwater would also


have led to a rapid sea-level rise, flooding continental shelves. As the ice retreated, therock debris and “flour” produced by mechanical ice weathering would have becomeexposed and commenced weathering. This likely led to an enhanced runoff of Ca2+

and bicarbonate, causing the shallow waters to become supersaturated with respect tocalcium carbonate and thereby enabling formation of more cap carbonates. As deep-water circulation was reestablished and the deepwater oxygenated, dissolved Fe2+ wasoxidized, leading to the deposition of the banded iron formations. Because reducedmetals, such as Fe2+, inhibit precipitation of CaCO3, removal of the metals wouldhave facilitated deposition of more cap carbonates providing more carbon drawdownpower.

Given the slow pace of the geological carbon cycle, the continental weatheringreactions must have taken at least 104 to 106 y to lower atmospheric CO2 back to asteady-state level. This suggests that the end of a Snowball Earth episode was markedby a transient period of high PCO2

in which ultra-greenhouse conditions were present.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 methanewould have been oxidized fairly rapidly to CO2 via reaction with atmospheric O2.

25.4.3 Phanerozoic Innovations to the GlobalCarbon Cycle

Oscillations in the PCO2of the atmosphere continued into the Phanerozoic as shown in

Figure W25.2. A continuing correlation with global temperatures suggests that CO2 wasan important agent in climate control. Several glacial events have occurred during thePhanerozoic, most notably at 300 mybp and throughout the past 1.8 my. All have beenassociated with CO2 levels below 1000 ppm with one exception, a short-lived glaciation450 mybp during which PCO2

was approximately 5000 ppm. High-resolution reconstruc-tions of paleotemperatures and atmospheric PCO2

have demonstrated a pervasive, tightcorrelation. The record for the past 650,000 y has been obtained from analysis of gastrapped in ice cores retrieved from subpolar and polar latitudes (Figure W25.3). Unlikethe Snowball Earth episodes, the glaciations of the Phanerozoic have been restricted tothe mid- to high latitudes.

Three processes seem to have contributed to the lack of Snowball Earths during thePhanerozoic. First, solar luminosity has continued to increase. Second, after the breakupof Rodinia 830 mybp, subsequent supercontinent cycles have redistributed land massesso that they are no longer concentrated at low latitudes. Third, biological innovationssuch as evolution of biomineralization and land plants have led to enhanced storage ofcarbon 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 diversityof metazoans and eukaryotic phytoplankton at the end of the Proterozoic and begin-ning of the Phanerozoic (Cambrian period). This is attributed to two side effects ofthe Snowball Earth episodes: mass extinction events and rapid weathering. The former


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Phanerozoic PCO2levels 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

Royer, D. L., R. A. Berner, I. P. Montaenez, 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 latterincreased the rate of macro- and micronutrient transport into the soils, freshwaters,and ocean. One of the important innovations of the Cambrian biological explosionwas the evolution of organisms capable of generating hard parts. Calcium minerals(carbonates and phosphates) rapidly emerged as a dominant building material usedby shallow-water benthic invertebrates for shells and skeletons. During the Cambrian,various types of microbes, including cyanobacteria, contributed to the formation ofextensive shallow-water calcified reef structures.

The rise of large vascular land plants during the Devonian period accelerated theremoval of CO2 from the atmosphere and the addition of O2. Land plants also acceler-ate chemical weathering rates, facilitating erosion of the continental land masses andincreasing the rate of nutrient delivery to the oceans. The increasingly oxic atmo-sphere of the Devonian (Figure W8.4) stimulated further evolution of animals, leadingto increased bioturbation of soil and sediments and lower burial rates of organic matter.

Pelagic calcifying plankton rose to ecological prominence during the Mesozoic erafollowing the largest of the mass extinction events of the Phanerozoic eon that occurredat the end of the Permian period (Figure W21.1). The environmental driving forcesfavoring the calcifiers are thought to have included: (1) high rates of nutrient supply dueto enhanced continental weathering rates and (2) extreme oversaturation in the surfacewaters with respect to calcite. Dinoflagellates and diatoms also came to prominence at


(Günz-Mendel)Aftonian

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FIGURE W25.3

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

human activities have driven PCO2levels 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 threegroups of plankton are now responsible for most of the export flux of POC to the deepocean and its sediments.

The final element of the modern oceanic carbon cycle was established once bio-genic calcite starting depositing onto the deep seafloor. This provided three importantstabilizing feedbacks to the global carbon cycle. First, the presence of sedimentaryCaCO3 stabilized deepwater carbonate concentrations, which provides buffering ofatmospheric PCO2

on time scales of millennia. Second, accumulation of biogenic calciteincreased the CaCO3 content of the sediments and, hence, enhanced the decarbonationcomponent of the geological carbon cycle. Third, as described in Chapter W15.6, shiftsin the CCD enable operation of a calcite compensation feedback that acts to stabilizeatmospheric PCO2

by altering the spatial extent of marine sediments in which CaCO3

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


always lie above the CCD. Hence, they cannot support a negative feedback loop thatbuffers against PCO2

change as effectively as can the deep-sea sediments. In the case ofthe latter, a decline in atmospheric PCO2

leads to a deepening of the CCD and, hence,an increase in the spatial extent of calcite deposition in the deep-sea sediments. As perEq. 25.1, calcite deposition serves as a net source of CO2 to the atmosphere. Becausethe shallow-water sediments already lie above the CCD, a decline in PCO2

serves onlyto increase the degree of calcite supersaturation in their overlying waters. If the super-saturation becomes high, calcite precipitation rates increase, but this effect is likelycountered by a concurrent reduction in the volume of the shallow waters and spatialextent of their underlying sediments. The latter is a consequence of lowered sea levelassociated with glaciation induced by global cooling during times of low PCO2

.To summarize, the biological innovations of the Phanerozoic have established a vari-

ety of feedbacks, some stabilizing and some destabilizing, that act on the atmosphericCO2 reservoir. Examples of how these feedbacks have operated during the major climateshifts of the Phanerozoic are described next following a short summary of how theglobal carbon cycle interacts with the atmospheric O2 and sulfur cycles.

Stabilization of the Atmospheric O2 by the Carbon andSulfur CyclesIn 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 linkagesamong the carbon, sulfur, and oxygen cycles. Sulfur is removed from the oceans viathe formation of sedimentary pyrite and the precipitation of CaSO4. The burial oforganosulfur compounds may also be important. Some of the sulfide fueling pyrite for-mation is hydrothermal in origin, but most comes from microbial sulfate reduction. Theprecipitation of CaSO4 occurs in hydrothermal vents and evaporites.

The reduced sedimentary sulfur acts as a significant sink of O2 when tectonic upliftexposes pyrite to the atmosphere. The exposed pyrite reacts with atmospheric O2. Thischemical weathering process oxidizes the pyrite’s sulfur and iron. Uplifted and exposedCaSO4 also undergoes chemical weathering in which congruent dissolution solubilizesdissolved sulfur as sulfate without affecting the atmospheric O2 reservoir. Any pyriteand CaSO4 that are not uplifted are carried back into the mantle via subduction. Duringthis process, the sulfur is converted into gaseous form (H2S and SO2) and emitted backinto the atmosphere via volcanic eruptions. Oxidation of this H2S and SO2 is anothersignificant sink of atmospheric O2.

Like the global carbon cycle, the sulfur cycle has undergone a massive reorganizationdue to the evolution of microbial life. In the case of sulfur, the two critical microbes arethe sulfide oxidizers and the sulfate reducers. Recall that the latter form a microbial con-sortium with the anaerobic methane oxidizers, thereby providing an important linkageto the carbon cycle (Chapter 7.3.2). The paleoceanographic record indicates that bothsulfate reducers and sulfide oxidizers evolved early in the Archean eon when the oceanwas still anoxic. In an ocean without O2, sulfide oxidation would have taken place viaanoxygenic photosynthesis in which sulfide (−II) is oxidized to elemental sulfur(0).


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 theatmosphere and its weathering (via aerobic respiration) is a net sink. Negative feedbacksinvolving sedimentary organic matter and pyrite serve to stabilize atmospheric PO2

. Forexample, increased PO2

promotes an enhanced oxidation rate of uplifted sedimentaryorganic carbon and pyrite, thereby consuming excess O2. This feedback operates only ifthe atmospheric and seawater levels of PO2

are rate limiting. In the case of the former,this requires that tectonic uplift and physical weathering be fast enough to keep upwith rising PO2

. In seawater, evidence for the controlling effect of O2 is provided bythe 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 sed-iments is dependent, at least in part, on O2 levels, which are in turn greatly influencedby the degree to which sulfur is partitioned among its reduced and oxidized reservoirs.The level of oxidation of seawater also influences CaCO3 precipitation rates as Mn2+

and Fe2+ inhibit the formation of calcite. This provides yet another linkage between thecarbon, sulfur, and O2 cycles.

25.4.4 Carbon Cycle Perturbations in the Phanerozoic

Some very large swings in PCO2, PO2

, and temperature have occurred during the Phanero-zoic. Considerable attention is focused on understanding the causes of these swingsand how the global carbon cycle responded in the hopes of obtaining insight into thelikely consequences of our anthropogenic additions to the atmospheric CO2 reservoir.Our additions have driven atmospheric PCO2

from its preindustrial level of 280 ppm to386 ppm (circa 2009).2 The last time PCO2

was this high was 34 to 40 mybp. Since thislast period of high PCO2

, Earth’s climate has shifted from an ice-free mode into onecharacterized by rapid cycling between glacial and interglacial conditions accompaniedby oscillations in global temperature and atmospheric PCO2

. Some scientists hypothesizethat our anthropogenic CO2 emissions will drive atmospheric PCO2

over a threshold thatwill switch Earth’s climate out of these “rapid” glacial-interglacial cycles and into someother mode that will likely start with a large-scale deglaciation. While PCO2

may not inand of itself initiate global climate change, its remarkable covariance with temperatureand 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 CycleAt present, paleoceanographers have identified four periods of significant climate changein the Phanerozoic. All were accompanied by major extinction events and, hence, havebeen 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, asreported by Dr. Pieter Tans, NOAA/ESRL (www.esrl.noaa.gov/gmd/ccgg/trends).


early Permian, 300 to 250 mybp. During the late Carboniferous, the longest and mostsevere glaciation of the entire Phanerozoic eon ended abruptly at 290 mybp. Within 10million years, PCO2

rose from present-day levels to 3500 ppm (Figure W25.2)! The amountof carbon injected into the atmosphere was equivalent to our entire modern-day fossilfuel reservoir. This was accompanied by a steep drop in PO2

, with levels decreasingfrom an all-time high of 25% v/v down to 16% v/v (Figure W8.4). The high PCO2

ledto global warming and deglaciation. Atmospheric PCO2

then underwent several short,high-amplitude oscillations in which intervals of low PCO2

supported transient periodsof glaciation. This oscillating finally ended around 270 mybp with PCO2

levels stabilizedsomewhere between 2500 and 3500 ppm, leaving Earth’s climate at a relatively warmstate that persisted until the end of the Mesozoic Era, 100 mybp. The beginning parts ofthis scenario are ominously similar to the current day, suggesting that we will eventuallysee 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 massextinction 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 secondbehind the end Permian extinctions. All three events coincide with periods of intenseflood basalt formation (Figure W25.4). Some of these massive piles of frozen lava are mil-lions 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 hotspot that is building the Hawaiian Island chain (Chapter 19.2.3). Geologists hypoth-esize that Earth’s crust is built by two cycles: one driven by plate tectonics, i.e., theSupercontinent 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 lattercontributes significantly to building of continental land masses.

The largest of these flood basalts, the Siberian traps, were laid down 250 mybp atthe Permian-Triassic-boundary. The Central Atlantic Magmatic province flood basalts,located in Central America, were deposited at the Triassic-Jurassic boundary and theDeccan traps, located in India, at the end of the Cretaceous. Exactly how these floodbasalts might have caused massive die-offs is not known but is thought to result froma combination of factors such as (1) emissions of volcanic CO2 and other acid volatilesresulting in acid rain and (2) ejection of huge amounts of ash into the atmosphere. Thelatter would have reflected insolation and thereby stopped or reduced photosynthesis.

The flood basalts were formed at times when large increases in PCO2occurred. In the

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 crisisat the end of the Triassic as acidification of the ocean was intense enough to cause thecollapse of carbonate reef platforms on the continental shelves. Global warming duringthese events was probably enhanced when rising temperatures eventually led to themelting of gas hydrates in the marine sediments, causing the release of methane.

The last large warming event in the Phanerozoic occurred at the Paleocene-Eoceneboundary and is called the Paleocene Eocene Thermal Maximum (PETM). The climateduring the early Eocene (52 to 55 mybp) was the warmest of the past 65 million years.


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end Jurassic?

end Triassic

end Pliensbachian

end Permian

end Guadalupian

Vilu

y (S

iber

ia)

end Frasnian ?

Em

elsh

anS

iber

ia

Cen

tral

Atla

ntic

Mag

mat

ic P

rovi

nce

karo

o an

d Fa

rrar

Par

ana

and

Ete

ndek

a

Raj

mah

al, K

ergu

elen

, Ont

ong

Java

(pha

se 1

)

Mad

agas

car

and

Car

ibbe

an P

late

au, O

nton

g Ja

va (p

hase

2)

Dec

can

Nor

th A

tlant

ic V

olca

nic

Pro

vinc

e (p

hase

1)

Eth

iopi

a an

d Ye

men

Col

umbi

a

100 150 200 250 300 350

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◦Cover a period of 10 to 20 ky. Atmospheric CO2 rose as high as 1000 to 1500 ppm. Thisclimate change caused rapid (10,000 y) changes in the species composition of the terres-trial flora. The PETM is thought to have been triggered by increased seafloor spreadingalong the Mid-Atlantic Ridge near Iceland. In addition to emitting volcanic CO2, thisenhanced 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. Thisled to warming of the deep waters and possibly the release of methane via melting ofthe gas hydrates buried in the marine sediments. The PETM is viewed as an importantanalog to present-day global warming as the rates and magnitudes of carbon release aresimilar to the present-day anthropogenic inputs.


80

0

500

1000

1500

2000

2500

3000

CO

2 (p

pm)

70 60 50 40 30 20 10

Time (Ma)

Cret. Paleogene

PaleosolsStomataPhytoplanktonBoron

Neogene

0

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 latestlarge-scale shift, transitioning into a relatively cold period that has persisted to date. Thisshift is thought to be driven by a decline in seafloor spreading rates and a concurrentincrease in weathering rates, both of which led to lowered CO2 (Figure W25.5). Thecooler 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 thelate Miocene, 10 mybp.) The formation of the Antarctic ice sheet is thought to haveincreased the rate of meridional circulation. Fast surface-water circulation rates andenhanced nutrient availability enabled the diatoms to come to ecological dominance.Because diatoms have a high export efficiency for POC, their rise caused the soft tissuepump to speed up, causing a further drawdown of CO2 and more cooling.

Fluctuations in Solar RadiationAt some point during the Oligocene, PCO2

levels declined to levels low enoughthat changes in insolation could significantly affect climate. This initiated the glacial-interglacial oscillations that have continued through to the present day with increase inamplitude in the Pleistocene (0 to 1.8 mybp). The insolation changes driving theseglacial-interglacial cycles are caused by shifts in Earth’s distance from the Sun. Theseshifts are driven by periodic variations in four orbital parameters: (1) the elliptical shapeof Earth’s orbit, (2) the tilt or obliquity of Earth’s axis relative to the plane of its orbit,


(3) the wobble, or precession, of Earth’s axis, and (4) shifts in the inclination of theplane of Earth’s orbit relative to the rest of the solar system. The periods of these orbitalvariations are on the order of 100,000, 41,000, 23,000, and 70,000 y, respectively, andare called Milankovitch cycles. At any given time, the amount of insolation receivedby the Earth can be predicted by considering the combined effects of the four orbitalparameters. Since the Oligocene, the Milankovitch cycle with the 100,000-y period hasdemonstrated the strongest correlation with global climate change, i.e., the timing ofglacial and interglacial conditions.

Other potential drivers of insolation changes that are large enough to affect climateinclude changes in atmospheric albedo and in solar irradiance. Various natural phenom-ena increase atmospheric albedo. These include volcanic emissions of gas and ash. Forexample, the volcanic gas, sulfur dioxide, is converted into sulfuric acid aerosols thatreflect radiation. The cooling effect of volcanic eruptions is shown in Figure 25.8a.Other important natural sources of reflective particles are: (1) aeolian dust transportedfrom deserts and semiarid regions and (2) smoke from forest fires. The intensity of thesesources is linked to climate.

Changes in solar irradiance are thought to occur over many time scales with themost well known being the sunspot cycle. During historical times, sunspot activity hasbeen observed to oscillate with an average period of 11 years. During the low part ofthe cycle, little-to-zero sunspot activity occurs. Interruptions to this cycle have involvedextended 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 witha 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

Thousands of years before present (B.P.)

18

24

22

0

2

16 14 12 10 8 6 4

Holocene epoch

Younger Dryas

Pleistocene epoch

Cha

nge

in te

mpe

ratu

re (8

C)

Little Ice age

Medieval warm periodHolocene maximum

Heinrich event 1

158C

2 0

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).


(1100 to 1250 AD), coincided (at least partially) with a period of warmth called theMedieval Warm Period. In more recent times, extended periods of lower sunspot activityhave taken place, namely from 1885 to 1910 and 1960 to 1975. Although these dropshave 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 hypoth-esize that positive feedbacks could amplify the effects of reduced solar irradiance onglobal climate, thereby providing a potential causal linkage between solar irradianceand global climate. One such proposed feedback involves a reduced photochemicalproduction of ozone in the stratosphere when solar irradiance is low. Since ozone is agreenhouse gas (GHG), its lowered production enhances the cooling effect of reducedsolar 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 nextMaunder-type minimum is not possible. On millennial time scales, evidence for changesin solar irradiance have been obtained from the polar ice record. Downcore variationsin the 10Be content of the ice are interpreted as a record of changes in solar irradiancebecause 10Be is produced in the atmosphere by chemical reactions whose productionrates are affected by solar irradiance. The ice core 10Be record goes back 9500 y andindicates that solar activity has been highly variable throughout the Holocene. Twomaxima in solar activity have occurred, one at 8000 ybp and the other 2000 ybp. Sincethat time, solar activity has generally been declining.

Ice Ages in the PleistoceneOver the past 1.8 million years, Earth has cycled in and out of Ice Ages during whichlarge continental ice sheets have waxed and waned. These changes in ice volume andregional 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 withwarm and interglacial conditions. Conversely, cold and glacial conditions have coincidedwith low GHG levels.

During the late Quaternary, Earth has been cycling in and out of Ice Ages every100,000 years or so, corresponding to the Milankovitch ellipticity orbital cycle (seeFigure 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 thatwe are in now. While the ice core data show us that some periods of rapid tempera-ture change were not accompanied by GHG changes, all periods of rapid and significantatmospheric PCO2

change have been accompanied by temperature swings. Similar behav-ior is seen in the other GHGs, CH4 and N2O. In the case of methane, atmosphericconcentrations are not as well correlated with climate shifts.

The role of the GHGs in climate change is still a matter of hot debate and researcheffort. While the Milankovitch cycles are the major agent of long-term climate change,


100 200Age (years) 3 103

300 400 500 6000

200

Nitr

ous

oxid

e (p

pb)

Met

hane

(ppb

)

220

240

260

280

300

400

800

1200

Methane

Carbon Dioxide

Nitrous Oxide

dD

1600

350

300

250

200

Car

bon

dixo

ide

(ppm

)Te

mpe

ratu

re p

roxy

(dD

‰)

2440

2420

2400

Marine Isotope Stage11 12 13 14 15 16

2380

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 climateresponds to changes in insolation. One of the key features of the atmospheric PCO2

,PCH4

, and temperature oscillations is their slow decline into glacial conditions followedby a rapid return to high levels during the interglacials. This suggests that a thresholdis reached that triggers a rapid shift from glacial to interglacial conditions. It also sug-gests that whatever drives Earth into an Ice Age is not necessarily the reverse of whathappens to end the Ice Age and bring Earth back to interglacial conditions.


Most scientists agree that oscillations in the GHG concentrations over Earth’s historyhave affected past climates through a variety of physical, chemical, biological, and geo-logical feedbacks, some positive and some negative. Various triggers and feedbacks thatare 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 oceanicprocesses. For example, high atmospheric dust loadings are a common characteristicof maximum glacial conditions (Figure 5.13). This suggests an important role of dustin (1) reflecting insolation and (2) enhancing the drawdown of atmospheric PCO2

. Thelatter 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 exportefficiency of POC.

Although the glacial-interglacial cycling in the late Quaternary is well correlatedwith oscillations in temperature and the GHGs, these oscillations have been modestcompared with those of earlier times. Atmospheric CO2 levels have varied by only 80to 100 ppmv and methane by 350 ppb, while temperature has shifted by only 10◦Cbetween the glacial and interglacial states. This suggests that the Earth’s carbon cycleand climate have been stabilized by powerful negative feedbacks that prevent theoccurrence of runaway ice houses or runaway hot houses as seen in earlier times. Themost 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 inenough of a climate change to shift Earth between glacial and interglacial states. Thissuggests that the human-induced changes in atmospheric PCO2

could have a substantialimpact on the timing of our planet’s next Ice Age. Based on the past 100,000-y period-icity of the glacial-interglacial cycling in the late Quaternary, Earth should be enteringor close to (within a few thousand years or so) another period of glaciation. But all betson this are now off, given that anthropogenic forcing has driven PCO2

to levels last seen30 million years ago during the PETM. As discussed in Chapter 25.5, humans are nowthe most likely drivers of global climate change.

No single process is likely responsible for causing Earth to shift between glacial andinterglacial states. Current hypotheses include a series of progressive steps over whichvarious positive feedbacks successively kick in. An example is given in Figure W25.8 inwhich intermediate conditions are envisioned as preceding full-blown glacial or inter-glacial states. The progression starts with an insolation change associated with one ormore of the Milankovitch cycles. The importance of this change is not in its overallimpact 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 meltingof ice that accumulated during winter. Changes in ice cover can lead to changes innutrient delivery to the ocean via terrestrial runoff and ocean circulation. These in turnaffect the marine biota. The combined physical and biotic changes will, in turn, alterthe rate at which the ocean’s carbon pumps can operate.

The scenario illustrated in Figure W25.8 starts with a physically driven descentinto glacial conditions. The ocean enables an initial PCO2

drawdown through two


Tab

leW

25.1

Oceanic

Pro

cesses

and

Their

Postu

late

dR

ole

sas

Triggers

and

Feed

backs

inG

lob

alC

limate

Change.

Po

stul

ated

Feed

bac

kD

escr

ipti

on

Oce

anic

Reg

ions

Infl

uenc

edM

echa

nism

s

BIO

TIC

PH

EN

OM

EN

A

Iro

nfe

rtiliza

tio

nIn

cre

ased

inp

uts

of

iro

n-r

ich

dust

incre

ase

marine

pro

ductivi

ty,

carb

on

up

take

insurf

ace

wate

rs,

and

sub

seq

uent

carb

on

flux

toth

e

deep

ocean,

whic

hlo

wers

atm

osp

heric

CO

2.

Reg

ions

influ

enced

by

aeo

lian

dust

inp

ut

and

with

und

eru

tiliz

ed

nutr

ient

(N,

P,

Si)

inve

nto

ries

(So

uth

ern

Ocean,

No

rth

Pacifi

c,

eq

uato

rial

Pacifi

c).

Who

le-o

cean

nutr

ient

incre

ase

Incre

ased

ocean

nutr

ient

reserv

oir

alle

viate

s

nutr

ient

limitatio

nand

allo

ws

glo

bal

incre

ase

inm

arine

bio

mass

and

carb

on

exp

ort

to

deep

ocean.

Nutr

ient-

limited

reg

ions

(e.g

.,lo

w-latitu

de

gyr

es).

Nutr

ient

utiliz

atio

nIn

cre

ased

utiliz

atio

no

fcarb

on

and

nutr

ients

insurf

ace

wate

rsre

mo

ves

CO

2fr

om

co

n-

tact

with

the

atm

osp

here

and

results

inlo

wer

atm

osp

heric

CO

2.

Incre

ased

nutr

ient

utiliz

a-

tio

nco

uld

occur

either

by

incre

asin

gsurf

ace

pro

ductio

no

rb

yd

ecre

asin

gve

rtic

al

mix

ing

and

there

fore

the

sup

ply

of

nutr

ients

and

CO

2to

surf

ace

wate

rs.

Reg

ions

with

und

eru

tiliz

ed

nutr

ient

inve

nto

-

ries

(prim

arily

the

So

uth

ern

Ocean,

eq

uat-

orial

Pacifi

cand

No

rth

Pacifi

c,

with

the

So

uth

ern

Ocean

co

nta

inin

ga

majo

rity

of

unused

nutr

ients

havi

ng

astr

ong

co

nnectio

n

with

CO

2-r

ich

deep

wate

rs).

PIC

-to

-PO

Cra

inra

tio

Decre

ased

exp

ort

ofC

aC

O3

rela

tive

too

rganic

carb

on

fro

msurf

ace

tod

eep

ocean

main

-

tain

shig

hsurf

ace-w

ate

ralk

alin

ity

(makin

gC

O2

mo

reso

lub

le)

and

red

ucin

gd

eep

-ocean

carb

onate

burial.

Asso

cia

ted

who

le-o

cean

alk

alin

ity

incre

ases,

red

ucin

gatm

osp

heric

CO

2.

Main

lyin

reg

ions

do

min

ate

db

yco

cco

lith-

op

ho

rid

pro

ductio

nto

day

but

ag

lob

al

shift

exp

ecte

d.

Silica

leakag

eM

echanis

mb

yw

hic

hth

era

inra

tio

co

uld

chang

e.

Iro

n-f

ert

ilize

dd

iato

ms

inth

eS

outh

ern

Ocean

take

up

less

silicic

acid

rela

tive

tonitra

te.

The

unused

silicic

acid

isexp

ort

ed

Reg

ions

no

rth

of

the

Anta

rctic

Po

lar

Fro

nt

that

are

silico

nlim

ited

.Lo

wla

titu

des

influ

-

enced

by

mix

ing

of

sub

anta

rctic

mo

de

wate

rin

toth

eth

erm

oclin

e.


tosilico

n-lim

ited

reg

ions,

incre

asin

g

dia

tom

pro

ductio

nat

the

exp

ense

of

co

cco

litho

pho

rid

s,

and

decre

asin

gth

e

CaC

O3-t

o-o

rganic

carb

on

rain

ratio

.

PH

YS

ICA

LP

HE

NO

ME

NA

Tem

pera

ture

and

salin

ity

chang

es

Mila

nko

vitc

hcyc

les

alter

inso

latio

nle

ad

ing

to

tem

pera

ture

chang

es.

Chang

es

inic

evo

lum

e

alter

salin

ity.

The

effect

of

tem

pera

ture

on

CO

2

so

lub

ility

iso

pp

osite

toth

at

of

salin

ity,

there

by

co

unte

ring

so

me

of

the

tem

pera

ture

effect.

Who

leo

cean.

Shelf

sed

iments

As

sea

leve

ld

eclin

es,

exp

osure

of

org

anic

-

rich

shelf

sed

iments

tote

rrestr

ial

weath

ering

incre

ases

deliv

ery

of

nutr

ients

and

alk

alin

ity

toth

eo

cean.

As

sea

leve

lin

cre

ases,

this

deliv

ery

declin

es.

The

exp

ort

of

PO

Mand

PIC

reb

uries

nutr

ients

and

alk

alin

ity

into

the

shelf

sed

iments

.

Co

asta

lw

ate

rs.

Chang

ein

ocean

alk

alin

ity

Co

nsid

ers

just

the

chem

ical

resp

onse

in

whic

ha

chang

ein

alk

alin

ity,

such

as

fro

m

altere

dte

rrestr

ial

inp

ut,

affects

the

so

lub

ility

of

CO

2.

Such

chang

es

wo

uld

eve

ntu

ally

alter

pre

serv

atio

no

fcarb

onate

sed

iments

via

carb

onate

co

mp

ensatio

nas

per

Chap

ter

W1

5.6

.1.

Who

leo

cean.

Chang

es

inp

ola

rcircula

tio

nC

aused

by

chang

es

inte

mp

era

ture

,salin

ity,

and

sea

ice

co

vera

ge.

Mix

ing

rate

sin

po

lar

wate

rsco

ntr

ol

the

deg

ree

tow

hic

hC

O2

can

be

reta

ined

.Lo

wm

ixin

gd

ue

toenhanced

density

str

atific

atio

nin

cre

ases

CO

2re

tentio

n.

Itals

oalters

nutr

ient

transp

ort

.

Sub

po

lar

and

po

lar

oceans.

So

urc

e:

After

Ko

hfie

ld,

K.

E.,

C.

Le

Quere

,S

.P.

Harr

iso

n,

and

R.

F.A

nd

ers

on.

(20

05

).R

ole

of

marine

bio

log

yin

gla

cia

l-in

terg

lacia

lC

O2

cyc

les.

Scie

nce

30

8,

74

–7

8;

and

Peaco

ck,

S.,

E.

Lane,

and

J.

M.

Restr

eo

(20

06

).A

po

ssib

leseq

uence

of

eve

nts

for

the

genera

lized

gla

cia

l-in

terg

lacia

lcyc

le.

Glo

bal

Bio

geo

chem

ical

Cyc

les

20

,G

B2

01

0.


40

280

270

260

250

240

230

220

Atm

osph

eric

PC

O2

210

200

190

18020 60 80 100 120

InterglacialIntermediateGlacialInterglacial

Increase in Southern Ocean temperature

Increase in southern Ocean mixing

Increase in low-latitude and northern temperature

Decrease in mean-ocean alkalinity and phosphate

Increase in northern hemisphere mixing

Reduction in temperature

Increase in mean-ocean alkalinity and phosphate

Fall in sea-level; rise in mean-ocean salinity

Reduction in high-latitudemixing

Rise in sea-level; decrease in mean-ocean salinity

Time (thousands of years)

FIGURE W25.8

Atmospheric PCO2over 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) reducedocean 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 sedi-ments to terrestrial weathering, thereby increasing nutrient and alkalinity delivery tothe ocean. This stimulates the marine biota to produce more POC and PIC, both ofwhich are then exported to the deep sea and sediments. Burial of POC and PIC fur-ther enhances the PCO2

drawdown and Earth’s descent into a full-blown glacial state.The return to interglacial conditions is envisioned as starting with a Milankovitch-drivenincrease in temperature that increases mixing in the Southern Ocean, thereby reducingits 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 andwarming act to reduce the rate at which the oceanic carbon pumps operate, caus-ing a continuing rise in atmospheric PCO2

. This provides a positive feedback to globalwarming that continues until full-blown interglacial conditions are attained.

Likely important details in these scenarios include shifts in plankton communitystructure induced by changes in nutrient and micronutrient availability. Such shifts


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 climateand ocean circulation. Climate controls the aeolian transport of iron. Ocean circulationcontrols 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 micronutri-ent bioavailable. The speciation calculations presented in Chapter 5.7 are illustrative ofefforts now underway to predict how much bioavailable iron can be generated from theparticulate forms. Because iron reduces the silica requirement of diatoms, iron fertiliza-tion of the Southern Ocean could be a key agent of control for the biological carbonpumps (Chapter 16.6.2). Plankton are also important climate control agents as they arethe 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 eitherdrive, or are driven by, meridional overturning circulation. For example, changes in therates of deepwater formation and its return to the sea surface should affect the rateat which nutrients are returned to the euphotic zone, hence altering the export fluxof PIC and POC. For example, a major slowdown in meridional overturning circulationis posited to have occurred 11,500 to 12,900 ybp as a result of a significant coolingevent (about 2◦C) called the Younger Dryas (Figure W25.6). This cooling was a briefinterruption of the general warming that has been ongoing since the last Ice Age ended18,000 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 NorthAtlantic. As Earth warmed, glacial meltwaters pooled in large lakes located at the presentsite of Northern Canada. Once the ice retreated sufficiently northward, the meltwaterswere free to flow into the North Atlantic. The flow path seems to have been acrossthe Great Lakes, down the St. Lawrence River, and into the Labrador Sea, one of thepresent-day sites of NADW formation. The presence of large amounts of freshwatershould have reduced surface-water salinities. Since the temperature at which seawatercan attain its highest density is inversely related to its salinity, fresher waters cannotget 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 fortwo reasons. First the heat that would have otherwise been released to the atmosphereby 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 Streamfrom reaching subpolar latitudes. Reestablishment of meridional overturning circulationtook place 1000 y later, probably as a result of cooling that enabled regrowth of the icesheets, thereby halting the flow of meltwater to the North Atlantic. The cause of thiscooling is unknown.

The formation of polar ice in the Northern Atlantic is also considered an importantcontrol of meridional overturning circulation, making it another likely driver of rapidclimate change in the late Quaternary. This linkage is of particular concern as the rateof loss of polar ice in the northern hemisphere is accelerating.


The power of meridional overturning circulation in altering global climate derivesfrom 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 similarpotential. They include: (1) the lateral spread of meltwaters on the sea surface in theform of eddies because of their potential to disrupt how the geostrophic currents trans-port heat and (2) changes in tropical rainfall rates as this affects the degree to whichsurface salty water can be returned from the surface waters of the Pacific and IndianOceans back to the site of NADW formation in the North Atlantic.

Warming in the HoloceneThe cooling event of the Younger Dryas (10,000 ybp) was the last gasp of the Wurmglaciation. The climate changes that have ensued during the continued transition intoan interglacial condition are illustrated in Figure W25.6. The step shift into warmertemperatures around 10,000 ybp defines end of the Pleistocene and the beginning ofthe Holocene epoch. As in past transitions from a glacial to an interglacial state, warm-ing has proceeded at varying rates. The major features have been a period of rapidwarming from 9000 to 5000 ybp called the Holocene Climatic Optimum during whichtemperatures were 0.5 to 2◦C warmer than in the present day. This was followed bya period of minor glaciation from 5000 to 2000 ybp that was in turn followed by twoshort and small warming events. The one that lasted from 900 to 1300 AD is called theMedieval warm period. A similarly short and small cooling event, called the Little IceAge, occurred from 1300 to 1850 AD. Since that time, global temperatures have beenrising (Figure 25.8). During these low-amplitude temperature shifts, atmospheric GHGlevels 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 thatoccurred prior to 1850. This impact is postulated to be a possible result of land-usechange, including deforestation and plowing, and biomass burning (forest fires). It ismore likely that the pre-1850 rise in CO2 was primarily due to “natural” processes inwhich the ocean’s calcite compensation played an important role as per the followingscenario. Warming at end of the last Ice Age melted continental ice, enabling regrowthof forests and wetlands. Most of the carbon fueling this regrowth was provided by thedegassing of oceanic CO2 in response to a dampened solubility pump. This degassing ledto a rise in carbonate ion concentrations in seawater by driving the following reactiontoward the products:

2HCO−3 → CO2−

3 + CO2 + H2O (W25.1)

High carbonate ion concentrations moved the CCD to deeper depths, leading toan increased rate of CaCO3 burial. This enabled release of even more CO2 from theocean to the atmosphere. Once the terrestrial biomass stabilized and atmospheric draw-down of CO2 ended, the ocean was no longer driven to supply CO2 to the atmosphereand the CCD returned to shallower depths.

Another interesting feature of the increase in PCO2over the Holocene is that the rise

has been sustained over a longer period than seen in the past three interglacials. Asshown in Figure W25.3, these interglacials have been characterized by an initial short


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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 highCO2 occurred 400,000 ybp. A possible explanation for this lies in the 400,000-y periodic-ity of the combined effects of some of the Milankovitch cycles. On this time scale, someof the orbital parameters are out of phase. This results in a combined effect on inso-lation that is smaller than usually seen during the typical 100,000-y glacial-interglacialcycle. Figure W25.3 also shows that first two interglacials in the late Pleistocene did nothave a well-defined initial increase in CO2 in contrast to the past five interglacials. Thisswitch in system response is attributed to changes in oceanic circulation and sea-surfacetemperatures in the Southern Ocean.

The last Ice Age terminated 10,000 to 15,000 ybp and has been followed by a periodof general warming termed the Holocene Epoch. Prior to 1700, atmospheric CO2 levelsduring the Holocene varied by no more than 10 ppmv, fluctuating between 275 ppmand 285 ppm. The increase in PCO2

that has taken place from 1700 to 2000 has been


about 85 ppm, equivalent to 175 Pg C, or 30% of the preindustrial level. AtmosphericCO2 concentrations reconstructed from ice cores extending back 650,000 y have estab-lished that in the Late Quaternary, the typical glacial-interglacial difference has been100 ppm, and was accompanied by temperature swings of about 10◦C (Figure W25.3).Thus the present-day PCO2

level, which now exceeds 380 ppm, represents a large depar-ture from that of the past 650,000 y (Figures W25.3 and W25.9). Most of this increaseis attributed to human activities as our known emissions are more than adequate toexplain the observed atmospheric increase (Figure 25.5). As shown in Figures 25.4 andW25.9, the time frame over which atmospheric PCO2

has risen rapidly coincides withthe period over which anthropogenic emission rates have increased. Figure 25.4b alsodocuments that the changes in PCO2

have been largest in the northern hemisphere,

matching the geographic source of our emissions. During this same period, the 14Ccontent of the atmospheric CO2 has decreased, further supporting a fossil-fuel source.

Figure W25.9 documents that the concentrations of other GHGs have also increasedover the past century. Methane levels have roughly doubled. Although still in the ppbrange, methane’s higher GWP makes its climate-forcing effect equivalent to 30% ofthat currently contributed by CO2. Fortunately methane is prone to oxidation, makingits atmospheric lifetime relatively short (8 to 10 y). The increase in its concentrationover 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 havenow stabilized.3 More than 50% of present-day global methane emissions are anthro-pogenic in origin, coming from fossil-fuel production (escape of natural gas from leakypipelines), livestock (cattle), rice cultivation, and waste handling (including animalwaste, domestic sewage, and landfills). Biomass burning has also been an importantsource, particularly from 0 to 1000 AD. Natural emissions of methane are primarily fromwetlands, peatlands, and tropical rain forests and from natural gas seeps associated withfossil-fuel deposits and the melt of gas hydrates. Swings in atmospheric methane lev-els during the end of the last glaciation are thought to be due to changes in wetlandproduction rather than methane-hydrate destabilization.

3 This stabilization is thought to have been caused by the drying of wetlands, leading to decreasednatural emissions and thereby countering some of the anthropogenic input. In 2007, PCH4

rose for thefirst time since 1998, jumping from an annual global mean of 1775 to 1783 ppb. Likely sources areincreased anthropogenic emissions from rapidly industrializing Asia and melting permafrost in the Arctic.

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