Art 1 Question 5 The Onset of the Glacial Periods Is More Abrupt Than the Interglacial Period

• Journal List
• Proc Natl Acad Sci U S A
• v.113(13); 2016 Mar 29
• PMC4822573

Proc Natl Acad Sci U Southward A. 2016 Mar 29; 113(thirteen): 3465–3470.

From the Encompass

Earth, Atmospheric, and Planetary Sciences

Carbon isotopes characterize rapid changes in atmospheric carbon dioxide during the concluding deglaciation

Thomas 1000. Bauska,^{a, b, 1} Daniel Baggenstos,^c Edward J. Beck,^a Alan C. Mix,^a Shaun A. Marcott,^d Vasilii 5. Petrenko,^e Hinrich Schaefer,^f Jeffrey P. Severinghaus,^c and James East. Lee^a

Thomas K. Bauska

^aCollege of Earth, Ocean, and Atmospheric Sciences, Oregon State University, Corvallis, OR, 97331;

^bSection of World Sciences, University of Cambridge, Cambridge CB2 3EQ, Britain;

Daniel Baggenstos

^cScripps Institution of Oceanography, University of California, San Diego, La Jolla, CA, 92093;

Edward J. Brook

^aCollege of Earth, Ocean, and Atmospheric Sciences, Oregon Country University, Corvallis, OR, 97331;

Alan C. Mix

^aCollege of World, Bounding main, and Atmospheric Sciences, Oregon State University, Corvallis, OR, 97331;

Shaun A. Marcott

^dDepartment of Geoscience, University of Wisconsin—Madison, Madison, WI, 53706;

Vasilii V. Petrenko

^eDepartment of Earth and Environmental Sciences, University of Rochester, Rochester, NY, 14627;

Hinrich Schaefer

^fClimate and Atmosphere Centre, National Institute of Water and Atmospheric Research Ltd, Wellington, New Zealand, 6023

Jeffrey P. Severinghaus

^cScripps Establishment of Oceanography, University of California, San Diego, La Jolla, CA, 92093;

James E. Lee

^aHigher of Globe, Sea, and Atmospheric Sciences, Oregon State Academy, Corvallis, OR, 97331;

Significance

Antarctic ice cores provide a precise, well-dated history of increasing atmospheric CO₂ during the last glacial to interglacial transition. However, the mechanisms that drive the increase remain unclear. Here we reconstruct a key indicator of the sources of atmospheric CO₂ by measuring the stable isotopic composition of CO_two in samples spanning the menstruation from 22,000 to 11,000 years ago from Taylor Glacier, Antarctica. Improvements in precision and resolution allow us to fingerprint CO_two sources on the centennial calibration. The data reveal two intervals of rapid CO₂ rise that are plausibly driven by sources from country carbon (at sixteen.three and 12.9 ka) and 2 others that announced fundamentally dissimilar and probable reflect a combination of sources (at 14.half-dozen and 11.5 ka).

Keywords: ice cores, paleoclimate, carbon bicycle, atmospheric CO₂, last deglaciation

Abstract

An agreement of the mechanisms that control CO₂ change during glacial–interglacial cycles remains elusive. Here nosotros help to constrain irresolute sources with a high-precision, high-resolution deglacial record of the stable isotopic composition of carbon in CO₂ (δ¹³C-CO₂) in air extracted from ice samples from Taylor Glacier, Antarctica. During the initial ascent in atmospheric CO₂ from 17.6 to xv.five ka, these data demarcate a subtract in δ^xiiiC-CO₂, probable due to a weakened oceanic biological pump. From 15.5 to 11.5 ka, the continued atmospheric CO₂ ascent of twoscore ppm is associated with small changes in δ^thirteenC-CO₂, consequent with a nearly equal contribution from a further weakening of the biological pump and rising ocean temperature. These two trends, related to marine sources, are punctuated at 16.3 and 12.9 ka with abrupt, century-calibration perturbations in δ^xiiiC-CO_two that suggest rapid oxidation of organic land carbon or enhanced air–sea gas exchange in the Southern Ocean. Boosted century-scale increases in atmospheric CO_ii ancillary with increases in atmospheric CH₄ and Northern Hemisphere temperature at the onset of the Bølling (xiv.6–14.3 ka) and Holocene (11.6–11.4 ka) intervals are associated with pocket-size changes in δ¹³C-CO₂, suggesting a combination of sources that included ascension surface ocean temperature.

Over 30 years agone ice cores provided the beginning articulate show that atmospheric CO₂ increased by about 75 ppm as Earth transitioned from a glacial to an interglacial land (i, 2). Afterward decades of research, the underlying mechanisms that drive glacial–interglacial CO₂ cycles are withal unclear. A tentative consensus has formed that the deglaciation is characterized past a internet transfer of carbon from the ocean to the atmosphere and terrestrial biosphere, through a combination of changes in ocean temperature, food utilization, apportionment, and alkalinity. Partitioning these changes in terms of magnitude and timing is challenging. Estimates of the glacial–interglacial carbon cycle budget are highly uncertain, ranging from twenty–30 ppm for the effect of ascension ocean temperature, v–55 ppm for sea apportionment changes, and 5–30 ppm for decreasing iron fertilization (3, 4), with feedbacks from CaCO₃ bounty accounting for upwardly to xxx ppm (five, six).

A precise history of the stable isotopic composition of atmospheric carbon dioxide (δ^thirteenC-CO_ii) can constrain central processes controlling atmospheric CO₂ (7, 8). A low-resolution tape from the Taylor Dome ice core (ix) identified a decrease in δ^xiiiC-CO₂ at the onset of the deglacial CO₂ rise that was followed past increases in both CO₂ and δ¹³C-CO₂ (Fig. 1). A college-resolution record from the European Project for Ice Coring in Antarctica Dome C (EDC) ice core (x) provided additional support for the rapid δ^xiiiC-CO₂ decrease associated with the initial CO_ii ascension, and box modeling indicated that this decrease was consequent with changes in marine productivity. The record likewise included other rapid changes in δ^xiiiC-CO₂, admitting at low precision, supporting large variations of organic carbon fluxes, notably a sharp increment in δ^thirteenC-CO_ii during the Bølling–Allerød (BA) interval attributed to carbon uptake by the terrestrial biosphere. A combined record including higher-precision EDC and Talos Dome data (11) documented a δ¹³C-CO₂ decrease beginning near 17.v ka. This shift in δ¹³C-CO₂ was interpreted to indicate that some process in the Antarctic ocean (Then), perchance changes in upwelling, drove the initial CO_ii ascension. This previous work did non resolve high-frequency variability in the δ^thirteenC-CO₂ records that may exist essential for discerning mechanisms of modify.

Carbon isotope records during the last deglaciation. Taylor Glacier δ¹³C-CO₂ information from this study (red). Previous work from Taylor Dome (gray open up circles) (9), Grenoble EDC data (open up dark-green squares) (10), Bern EDC data (orange circles) (11, 45), sublimation measurements from EDC (blue triangles), and Talos Dome (regal squares) with an estimate of the 1-sigma dubiousness from a compilation of previous ice cadre δ¹³C-CO₂ data (eleven).

Here we use an analytical method (12) that employs dual-inlet isotope ratio mass-spectrometry to obtain precision budgeted that of modern atmospheric measurements [∼0.02‰ 1-sigma pooled SD based on replicate analysis compared with ∼0.05–0.11‰ for previous studies (9–11)]. Nosotros extracted atmospheric gases from big (400–500 g) samples taken from surface outcrops of ancient ice at Taylor Glacier, Antarctica, at an average temporal resolution of 165 y between 20 and 10 ka, and subcentury resolution during rapid change events. This resolution allows us to delineate isotopic fingerprints of rapid shifts in CO_two that were previously impossible to resolve. Our study complements recent precise observations of CO₂ concentration variations during the last deglaciation, which revealed abrupt centennial-calibration changes (thirteen) (Fig. 2).

Carbon cycle changes of the last deglaciation. WAIS Split continuous CH_four (dark-green) (xiv) and detached CO_ii (blue) (xiii) concentration data plotted with Taylor Glacier CO₂ and δ¹³C-CO₂ data (this written report) (red markers, black line is a smoothing spline), the v-point running Keeling intercept with shading indicating the R² for each time interval. Blue bars indicate intervals of rapid CO₂ rise identified in the WAIS Divide ice core (thirteen).

During the initial 35-ppm CO₂ rise from 17.half-dozen to fifteen.5 ka, nosotros find a 0.3‰ decrease in δ¹³C-CO₂ that is interrupted by a sharp minimum ancillary with rapid increases in CO₂ and CH₄ around xvi.three ka (13, xiv) (Fig. 2). The xvi.3-ka characteristic in the CO₂ and CH₄ concentration records, which corresponds to a 0.1‰ negative excursion in δ¹³C-CO₂, has been plausibly tied to the timing of Heinrich effect 1 (13, fourteen) and signals a fashion switch in the deglacial CO_ii rise. The subsequent slower rise in CO₂ from 15.5 to fourteen.8 ka is not accompanied by large changes in δ¹³C-CO₂. Across the Oldest Dryas to Bølling transition (14.6–fourteen.3 ka) and coincident with a 10-ppm CO_ii increase and large CH₄ increase, nosotros resolve a 0.08‰ increase in δ^thirteenC-CO₂ (Fig. 2). Rapid increases in CO₂ and CH_four at the Younger Dryas (YD) to Preboreal transition (11.half dozen–11.four ka) are associated with minor variability δ¹³C-CO₂. On the other hand, the onset of the YD (12.8–12.5 ka) is characterized by a small ascent in CO_two associated with a 0.15‰ decrease in δ¹³C-CO₂ that appears tightly coupled to the timing of the large CH₄ decrease. The recovery from this excursion is characterized past increasing CO_ii and δ¹³C-CO_ii. Broadly, our information confirm the results of Schmitt et al. (11) (Fig. ane). However, some of the big swings in δ¹³C-CO₂ indicated by the earlier EDC record (10), may be inaccurate and crave reexamination.

Processes Controlling Atmospheric δ¹³C-CO₂

To visualize the constraints provided past δ^thirteenC-CO₂ on the processes controlling CO_ii nosotros use a cantankerous-plot of CO₂ and δ¹³C-CO₂ (referred to every bit a Keeling plot when the x axis is equal to 1/CO₂). When the archetype Keeling plot is applied to a ii-component system, and bold conservation of mass, the y axis intercept of a linear regression to the data (y ₀) approximates the δ¹³C signature of a secondary external source mixing with a primary source (15). In the more complex mixing between the atmosphere, ocean, and terrestrial biosphere, y ₀ is withal indicative of the source reservoir's δ^thirteenC signature merely interpretation requires a model of the carbon cycle to account for processes such as air–body of water gas exchange, ocean mixing, and ocean–sediment interactions, which buffer the atmospheric δ¹³C-CO_ii signature over long timescales and reduce the slope of δ¹³C:1/[CO₂] (seven). Nosotros use a simple box model of the carbon cycle, previously published box model experiments (7) and intermediate complexity models to business relationship for these effects and deconvolve the processes responsible for the deglacial ascent in CO_ii. We split the processes into the following categories: ocean productivity and circulation, land carbon storage (particularly rapid changes), ocean temperature, the CaCO₃ bicycle, and air–ocean gas substitution. Private model Keeling plot intercepts are listed in SI Appendix, Table S1; the ranges within each category are represented graphically in Fig. 3A .

Cross-plot of data constraints and model experiments. (A) Shaded lines show the range of model-based constraints on various carbon cycle processes as listed in SI Appendix, Table S1 (changes in sea biological pump/circulation, xanthous; deglacial increase in SST, blue; rapid release of land carbon, green; rapid modify in Southern Body of water gas exchange, majestic; CaCO₃ wheel, gray). (B) All Taylor Glacier data with arrows as guides to the approximate time path. (C) The information divided into the early HS1 (yellow) and afterwards deglaciation (blueish) modes of variability. Colored markers dissever the information by fourth dimension period and the shaded vectors indicate the linear regressions of the data with the 1-sigma doubtfulness. (D) Further partition of the data into the precipitous changes at The xvi.iii-ka event and onset of the YD (red). See SI Appendix, Table S2 for statistics.

An oceanic δ^thirteenC-DIC depth slope is controlled by a combination of ocean mixing and export from the surface of isotopically light organic carbon. Decreased carbon consign or increased ocean ventilation during the concluding deglaciation would decrease the δ^xiiiC-DIC gradient, leading to a ascent in atmospheric CO_two and decrease in atmospheric δ^thirteenC-CO_ii. First, we simulate a plausible signature of a glacial–interglacial weakening of the marine biological pump by forcing a decrease in the strength of the Subantarctic biological pump from near full efficiency (PO₄ = 0.2 mmol 1000⁻ⁱⁱⁱ) to preindustrial level (PO₄ = one.4 mmol grand^−three). The timing of this modify is scaled to the subtract in dust delivery to Antarctica (xvi). This leads to an increment in CO₂ and decrease of δ^xiiiC-CO_ii; the human relationship is characterized in our model with y ₀ equal to −eight.half dozen‰. For comparison, factorial experiments with the Bern 3D model of the last glacial–interglacial cycle isolate the effect of atomic number 26 fertilization in the model (17) with a resultant Keeling intercept of −9.vi‰. Second, nosotros vary the rate of SO upwelling in our box model. Greater ocean ventilation raises CO₂ and lowers δ¹³C-CO₂ resulting in a Keeling plot intercept of −eight.four‰. Experiments with the Bern 3D model where wind stress over the And so was varied (20–180%) evidence that atmospheric CO_two positively correlates with the rate of bounding main overturning (18). The simulated CO_ii and δ^thirteenC-CO₂ produce a Keeling plot intercept between −7.6‰ [when the Atlantic Meridional Overturning Circulation (AMOC) is in an "on" country] and −8.6‰ (when AMOC is in an "off" state). Finally, decreased AMOC has been hypothesized to slow the delivery of depression preformed food h2o to the deep ocean and consequently drive a weakening of the biological pump. Experiments with the Model of Ocean Biogeochemistry and Isotopes/University of Victoria climate model of intermediate complexity that simulated an interval of collapsed AMOC show atmospheric CO₂ ascension and δ^thirteenC-CO₂ decreasing (nineteen) (y ₀ = −eight.5‰). Combining all these experiments provides plausible constraints on oceanic sources to the atmosphere (y ₀ ∼= −seven.4 to −9.six‰), which include both productivity driven changes (east.g., iron fertilization) and deep-ocean ventilation modify (due east.g., enhanced turnover of deep water masses) (Fig. 3A ).

During the deglaciation, the release of oceanic carbon to the atmosphere is likely partially get-go by the gradual accumulation of carbon on land, primarily during the later part of the deglaciation when the major water ice sheets are small and dwindling (fourteen–ten ka) (20). This land carbon uptake would lower atmospheric CO₂ and increment δ¹³C-CO₂. Because the magnitude of carbon isotopic fractionation by photosynthesis is very like in the marine and terrestrial regimes, changes in organic carbon cycling between the land, atmosphere, and bounding main that are slower than the timescale of bounding main mixing are broadly indistinguishable from the atmospheric data alone. However, for changes in organic country carbon storage that are rapid relative to the mixing time of carbon in the body of water–atmosphere system, the atmospheric indicate will more than closely reverberate the isotopic signature of organic carbon and and then diminish as information technology is buffered by exchange with the deep ocean. We drive changes in atmospheric CO₂ of about 10 ppm by varying the land-to-atmosphere carbon flux in our model at periodicities of 50, 500, and v,000 y. The modify in δ¹³C-CO₂ decreases with increasing periodicity resulting in y ₀ of −thirteen.four‰, −10.9‰, and −9.8‰ at the 50, 500, and 5,000 y periodicities, respectively. Fast country carbon fluxes to the atmosphere can therefore be distinguished from changes in the body of water biological pump with high-resolution atmospheric data (y ₀ = ∼<−10.9‰; Fig. 3A ).

Increasing bounding main temperature decreases both the solubility of CO₂ in seawater and the magnitude of isotopic fractionation during air–sea gas substitution. Ascent atmospheric CO₂ and increasing δ¹³C-CO₂ are therefore consequent with ocean warming. Forcing our model with latitudinal temperature stacks (21) for the deglaciation results in a 35-ppm rise in atmospheric CO₂ with an increase of well-nigh 0.iii‰ in δ^thirteenC-CO₂ with an credible y ₀ from this effect of −iv.5‰ (Fig. 3A ).

Carbon isotopic fractionation during CaCO₃ formation from seawater is very pocket-size compared with that of photosynthesis, and the δ¹³C of CO₂ from volcanic emissions, though poorly constrained, is very similar to atmospheric values. Processes like CaCO₃ compensation, reef edifice, and volcanic emissions are thus consistent with rising CO₂ and lilliputian to no change in δ^thirteenC-CO_ii (y ₀ = initial atmospheric δ¹³C-CO_ii; Fig. 3A ). Moreover, a decrease in the corporeality of respired carbon in the deep bounding main, driven by either oceanic changes or land carbon regrowth, will trigger increases in CaCO₃ preservation (and corresponding production of CO₂) that act to restore [ CO 3 two − ] over multimillennial timescales (22, 23). The indirect effect of a weakened biological pump or state carbon regrowth would thus be an increase in CO_two with little change in δ^thirteenC-CO_two over thousands to tens of thousands of years. In Keeling plot space the inferred intercept of a weakened biological pump would slowly asymptote to the CaCO₃ intercept. On even longer timescales the δ¹³C-CO₂ signature of all processes is dampened toward a steady country determined by the input of volcanic and weathering fluxes of carbon to the temper/ocean (10⁵ y) (24). The box model experiments presented here largely exclude CaCO₃ feedbacks (SI Appendix) and thus represent only the direct effect of diverse carbon bicycle processes that are of import in constraining the isotope signature on the centennial-to-millennial timescale, at the expense of underestimating the long-term feedbacks that are significant on glacial–interglacial timescales.

Changes in air–sea gas exchange, via changes in sea-water ice extent or wind speed in the And so, are hypothesized to take a meaning bear upon on CO_two and δ^thirteenC-CO_ii, generally with increased air–sea gas exchange leading to increases in atmospheric CO_two and big decreases in δ^xiiiC-CO₂ (7, 25). In our model, varying the air–ocean gas commutation coefficient over the So past ±50% over periods of fifty, 500, and 5,000 y (but keeping ocean mixing constant) produces minor rises in CO₂ and a sharp decrease in δ¹³C-CO₂ when air–sea gas exchange is enhanced. Keeling plot intercepts vary greatly with the periodicity of forcing (−37‰ to −18‰) merely are consistent with results from the Box model of the Isotopic Carbon cYCLE (BICYLE) (SI Appendix, Table S1). Changes in air–body of water gas commutation could therefore contribute to rapid δ¹³C-CO₂ variability.

Identifying and Diagnosing Deglacial δ^thirteenC-CO₂ Variability

All of the higher up processes tin work in combination, leading to a arrangement that is fundamentally underconstrained past the CO_two-to-δ^xiiiC-CO₂ relationship. Nonetheless, the changing relationship between CO₂ and δ^thirteenC-CO_two with time can be combined with the model constraints to dissever the information into fourth dimension intervals based on the dominant processes (Fig. threeB ). We identify iv major patterns of variability of the carbon cycle spanning the length of our tape: (i) a millennial-scale increase in CO_ii and decrease in δ^xiiiC-CO₂ during the early office of Heinrich Stadial i (17.6–15.5 ka); (ii) ascent CO_two with generally increasing δ¹³C-CO₂ during the later portion of Heinrich Stadial i (15.5–14.half-dozen ka) and later portion of the YD (12.8–11.5 ka); (3) ascension CO₂ with centennial-scale negative isotopic excursions at 16.iii and 12.8 ka; and (iv) centennial-calibration CO_ii rises with minor changes in δ¹³C-CO₂ at 14.6 and 11.5 ka. SI Appendix, Table S2 provides the Keeling plot intercepts for some of these intervals and Fig. 3 C and D prove the divisions graphically.

From 17.6 to 15.5 ka δ^thirteenC-CO_two decreases by about 0.iii‰ and CO₂ increases by about 35 ppm. This singled-out stage of the deglacial CO₂ rise was previously identified with less precise data and attributed to an increase in SO upwelling (11). Excluding the excursion effectually 16.3 ka (see below) the stiff human relationship between CO₂ and δ¹³C-CO₂ across this interval (R ^two = 0.97, y ₀ = −viii.6‰; Fig. iiiC ) is consistent with the bulk of the CO_two increase being driven by a weakening of the efficiency of the biological pump. The decrease in atmospheric δ¹³C-CO₂ is thus consequent with either increased ocean ventilation (eleven, 26, 27), an increment in the ocean preformed nutrient content driven by a decrease of North Atlantic Deep Water (NADW) germination (28), or a decrease in the consign of organic carbon to the deep ocean (29).

The timing of the δ^xiiiC-CO_two subtract coincides with the deglacial subtract in the dust flux over Antarctica and may lead the inferred maximum in SO upwelling. By 16.0 ka, the not-sea-table salt calcium flux at the Talos Dome ice core site (30) had decreased to near interglacial levels, whereas the And so opal flux recorded at 53.2°S, 5.i°Eastward (31) was even so increasing to values that peaked between 15.five and 14.v ka (Fig. 4A ). Our data thus support the hypothesis that a decrease in iron fertilization was important during the earliest stages of the last deglaciation (32). The magnitude of the direct effect of iron fertilization is partially constrained by the data, suggesting an upper limit of 35 ppm CO₂ modify from this mechanism. This is consistent with empirically derived estimates from the relationship between atmospheric CO₂ and Subantarctic productivity over multiple glacial–interglacial cycles (∼40 ppm) (33) and the coupling of CO_ii and ice core proxies for dust delivery during the last glacial–interglacial bike (≤ 40 ppm) (xxx). Even so, state-of-the-art biogeochemistry models simulate smaller glacial–interglacial changes due to iron fertilization between eight and 15 ppm (34–36). Model and empirical estimates could be reconciled if other mechanisms for lowering the efficiency of the biological pump (due east.g., ocean ventilation) are working in concert with atomic number 26 fertilization during this interval and business relationship for part of the 35 ppm increment in atmospheric CO₂. Possibly, a fast response of the carbon cycle to iron fertilization is superimposed on a slower modify driven by upwelling, resulting in the 2 singled-out rates of CO_ii rise during HS1.

Climate and carbon cycle changes during the terminal deglaciation. (A) Proxies for Greenland temperature (46) (purple), Due west Antarctic temperature (37, 47) (blue), Due east Asian precipitation (48) (green), grit commitment to Antarctica (30) (yellow), Southern Ocean upwelling (31) (blueish markers), global temperature relative to the early on Holocene (blue banding) (21), and the Taylor Glacier CO_two and δ^xiiiC-CO_ii data (red). The red confined betoken periods of rapid δ^xiiiC-CO₂ decreases; blue confined betoken rapid CO₂ increases with slight increases or fiddling change in δ¹³C-CO₂. B and C highlight the changes in temperature and atmospheric CO_two at the centennial scale.

From 15.v to 11 ka, atmospheric CO₂ increases by 40 ppm and δ^xiiiC-CO₂ gradually increases (y ₀ = −6.2‰; Fig. 3C ) with a plateau during the BA (∼14.6–13.0 ka) at values of 244 ± 2 ppm and −vi.63 ± 0.04‰ for CO₂ and δ^thirteenC-CO₂, respectively. Broadly, the big increase in CO_ii and pocket-sized overall modify in δ¹³C-CO_ii is consistent with the atmospheric CO₂ increase existence driven past changes in the CaCO₃ cycle, volcanic emissions, or concurrent changes in organic carbon bike and ocean temperature. About studies conclude that CaCO_three feedbacks account for up to 30 ppm of glacial–interglacial CO_two change (with an due east-folding timescale of ∼5,000 y; refs. 5, six). Though certainly significant in controlling the glacial–interglacial CO₂ and δ^xiiiC-CO₂ differences, these effects are too dull to explain the rapid increases in atmospheric CO₂ effectually 14.75 ka (∼10 ppm over 200 y) and 12.9–11.5 ka (∼30 ppm over 1,500 y). Moreover, a CO₂ increase driven by only the CaCO_three bicycle or volcanic emissions would produce no variability in δ¹³C-CO₂. Instead, changes in δ¹³C-CO₂ of about 0.1‰ on the centennial timescale advise that these changes are in part driven past a combination of ascent SST and ¹³C-depleted sources (i.e., ocean ventilation, land carbon). Allowing for relatively small contributions from CaCO₃ cycling and volcanic emissions, the δ¹³C-CO_ii data from fifteen.five to xi.0 ka are consequent with a roughly equal mix of sources from rising sea temperature and a weakened biological pump. The trend to more positive δ^thirteenC-CO₂ suggests that the temperature event was slightly greater (60 ± ten%, assuming two finish-fellow member mixing). Annotation that whatever sources of atmospheric CO₂ from the CaCO₃ cycle or volcanic emissions would decrease the inferred absolute changes of these two processes just take niggling effect on their relative contribution. The greater importance for temperature-driven changes in the after compared with the earlier part of the CO₂ rise is consistent with the increase in global surface temperature lagging the increase in atmospheric CO₂ (21).

A recent high-resolution record from the West Antarctic Ice Sheet (WAIS) Divide water ice core demonstrated that rapid increases in CO₂ of virtually 12 ppm at both the onset of the BA (14.6 ka) and terminate of the YD (11.5 ka) occurred exactly ancillary with abrupt increases in CH₄ and Northern Hemisphere (NH) temperature (xiii). Our data advise that the effect of ascension ocean surface temperature (SST) on atmospheric CO₂ may be well-nigh pronounced during these ii distinct intervals. Moreover, the WAIS Divide water ice core revealed that Antarctic temperature remained stable or even continued to warm until ∼200 y after the onset of NH warming (37) (Fig. iv B and C ). A global temperature reconstruction, though uncertain on centennial timescales, records temperature increases of ∼one and ∼0.5 °C at the onset of the BA and cease of the YD, respectively (21). At the onset of the BA, our records testify a 12-ppm increment in CO₂ and a 0.1‰ increment in δ^xiiiC-CO₂, consistent with SST dominating the atmospheric CO₂ upkeep. At the end of the YD, we observe very petty change in δ^xiiiC-CO₂ during a 10-ppm rising in atmospheric CO_2, suggesting a balanced contribution from ¹³C-depleted carbon sources and rising SST. This relationship between ocean warming and rising CO_two suggests an important positive climate–carbon feedback that may exist operating on the centennial timescale. These observations besides constrain hypotheses that organic carbon sources explain the atmospheric CO₂ increases associated with NH temperature rising. Thawing NH permafrost at the onset of the BA (38) or ocean "flushing" events tied to the resumption of AMOC (39) would need to be compensated by carbon sinks that are more depleted in ¹³C (i.due east., a steeper vector in the Keeling plot that leads to a cyberspace increase in CO₂ and slight increment in δ¹³C-CO₂) or accompanied past sources that enrich the atmosphere in ^thirteenC.

Two significant features in our record are the sharp century-scale minima in δ^xiiiC-CO_ii centered at 16.3 and 12.8 ka. These ii events are associated with meaning increases in atmospheric CO_two of about 7 ppm and very rapid decreases in δ¹³C-CO₂ of nearly 0.2‰ (Fig. threeD ). The higher resolution WAIS Split up tape (13) indicates that the atmospheric CO₂ increase during the sixteen.iii-ka event is greater in magnitude (∼12 ppm) and more rapid than our Taylor Glacier data resolves. The sharp CO₂ increase at 16.3 ka could plausibly be interpreted equally an consequence superimposed on a iii-kyr-long tendency of rising atmospheric CO₂ from most 17.6–14.75 ka, whereas the 12.9-ka circuit appears to occur near the beginning of a relatively rapid atmospheric CO_two increase from 12.9 to xi.5 ka. The 16.3-ka event is as well associated with a small CH_iv increment that has been attributed to a rapid increment in Southern Hemisphere (SH) methane sources, possibly associated with a southward excursion of the intertropical convergence zone (ITCZ) associated with Heinrich event one (xiv).

Our modeling experiments show that these features are consistent with a rapid release of terrestrial carbon to the atmosphere over a menses of a few hundred years or less (accounting for the smoothing effect of gas trapping in the firn; SI Appendix, Fig. S10). These two events occurred during intervals of very weak monsoon strength in the northern tropics and some of the coldest conditions in the high-latitude NH (Fig. 4). Model experiments suggest that colder and drier weather condition post-obit collapse of the AMOC tin can bulldoze decreases in land carbon stocks in the high-latitude NH (40). Although the global net rest in the model depends on the background climate and vegetation, net increases in atmospheric CO₂ occur under glacial conditions. Our data thus propose a possible link between tropical CH₄ production and high-latitude terrestrial carbon pools driven by centennial-scale cold periods and/or drought during the deglaciation.

Alternatively, the minima in δ^thirteenC-CO₂ are consistent with rapid CO₂ increases driven in part by periods of enhanced air–ocean gas exchange. Every bit shown earlier, δ¹³C-CO_ii can exist highly sensitive to changes in air–sea gas exchange. The precipitous drops in δ¹³C-CO_two may reflect intervals of enhanced air–sea gas substitution that, in combination with other ¹³C-depleted sources, drive increases in atmospheric CO₂.

Another possible scenario that can produce a Keeling plot intercept of less than the typical oceanic terminate fellow member involves a combination of a weakening biological pump source that is moderated past a smaller CO₂ sink from decreasing SST. The combination of the two vectors could produce an intercept that is more negative (<−8.6‰) than the biological pump signature alone. Although this scenario is unlikely for the 16.three-ka effect where we observe a significant and rapid increase in atmospheric CO_two with no big SST decreases, the δ¹³C-CO_ii subtract at the onset of the YD is associated with a strong winter cooling in the NH, specially in the Northward Atlantic. The subsequent recovery of δ¹³C-CO₂ following the minimum would probable require a source of atmospheric CO_two from increasing SST, maybe from a delayed warming in the SH.

Decision

Many possible scenarios can explain the development of atmospheric CO₂ and δ¹³C-CO₂ during the deglaciation. Narrowing the range of scenarios is possible if the observed changes in carbon bike can be consistently coupled to the climate history. We conclude by outlining one possible scenario that links our observations of centennial-scale features and the broader millennial-calibration changes within the context of the deglacial climate transition (Fig. ivA ). During the early part of HS1, the collapse of the AMOC (41) decreased heat ship to the N Atlantic. In response, big areas of the NH cooled and the SH warmed; possibly the ITCZ shifted southward and SH westerlies shifted southward or strengthened (42–44). We hypothesize that a shift of the westerlies off the SH continents and/or increased SH precipitation led to a precipitous decline in dust delivery over the Subantarctic bounding main, driving upward to 35 ppm of the CO₂ rise from nigh 17.half dozen–15.5 ka. The s migration of the ITCZ also led to drying in parts of the NH, possibly causing a reduction in organic land carbon, most notably around xvi.3 ka. Alternatively, or additionally, the changing SH westerlies reached a threshold around xvi.iii ka, in which air current speed over the SO increased, leading to enhanced air–ocean gas substitution and possibly greater upwelling.

During the later half of HS1 (xv.5–14.6 ka), dust deposition in the SO had fallen to interglacial levels and farther CO₂ ascent was driven mostly by warming ocean temperatures and an additional weakening of the biological pump, with a peak in SO upwelling or an extended interval of AMOC plummet as two possible mechanisms. During the YD (12.9–11.5 ka) most of the CO_two increase was driven by similar processes to the later half of HS1. However, the initial rise in CO_two during the YD could have been driven by either a 2nd loss of land carbon or renewed enhancement of SH westerlies. At the onset of the BA (14.half dozen ka) and end of the YD (eleven.v ka) significant warmings in the NH and connected warming around Antarctica likely contributed to the centennial-scale increases in atmospheric CO₂.

The δ^thirteenC-CO₂ record shows that the deglacial increment in atmospheric CO_ii occurred in a series of steps, each with a δ¹³C fingerprint that suggests that dissimilar mechanisms may have been triggered at various times during the deglacial transition. Early in the transition, the subtract in δ¹³C-CO₂ is consistent with (albeit not uniquely) a combination of atmospheric CO_two sources from respired organic carbon that exceeded sources from rising ocean temperature. This suggests that the initial trigger for the deglacial CO_two rise involved either an ocean circulation or bounding main biological process. Afterwards in the transition, the relatively stable δ^thirteenC-CO₂, punctuated by centennial-scale changes, suggests a combination of sources that could include changes in the CaCO₃ cycle or volcanic emissions, just nigh probable reflects a counterbalanced contribution of respired organic carbon and rise body of water temperature that strengthens and weakens over time. At least twice during the deglaciation a rapid release of ^xiiiC-depleted carbon to the atmosphere may take occurred over a few centuries, suggesting that abrupt and significant releases of CO₂ to the atmosphere may be common nonlinear features of Earth'south carbon cycle. Farther work on defining the isotopic signature of glacial–interglacial CO_ii mechanisms beyond a suite of carbon cycle models could yield a more precise agreement of CO₂ sources during the deglaciation.

Supplementary Material

Supplementary File

Acknowledgments

We thank Tanner Kuhl, Robb Kulin, and Paul Rose for assistance in the field and Fortunat Joos, Laurie Menviel, and Andreas Schmittner for sharing model output. We thank Mathis Hain and an anonymous reviewer for comments that improved the newspaper. We are grateful for technical back up from NSF-funded Ice Drilling Design and Operations (University of Wisconsin) and the OSU/CEOAS Stable Isotope Laboratory, in detail Andy Ross. This work was funded by NSF Grants ANT 0838936 (Oregon State Academy) and Pismire 0839031 (Scripps Institution of Oceanography). Further support came from the Marsden Fund Quango from New Zealand Government funding, administered by the Royal Club of New Zealand and NIWA under Climate and Atmosphere Research Programme CAAC1504 (to H.Southward.). T.M.B. was partially supported by the Comer Science and Education Foundation.

Footnotes

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