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A preliminary estimate of changing calcrete carbon storage on land since the Last Glacial Maximum
Global and Planetary Change 20 Ž1999. 243–256 www.elsevier.comrlocatergloblacha
A preliminary estimate of changing calcrete carbon storage on land since the Last Glacial Maximum J.M. Adams b
a,b,)
, W.M. Post
a
a MS 6335, EnÕironmental Sciences DiÕision, Building 1509, Oak Ridge National Laboratory, Oak Ridge, TN 37831, USA Department of EnÕironmental and Geographical Sciences, UniÕersity of Adelaide, Napier Building, Adelaide, SA 5005, Australia
Abstract The glacial-to-interglacial shift in land carbon storage is important in understanding the global carbon cycle and history of the climate system. While organic carbon storage on land appears to have been much less than present during the cold, dry glacial maximum, calcrete Žsoil carbonate. carbon storage would have been greater. Here we attempt a global estimation of this change; we use published figures for present soil carbonate by biome to estimate changing global soil carbonate storage, on the basis of reconstruction of vegetation areas for four timeslices since the Last Glacial Maximum. It appears that there would most likely have been around a 30–45% decrease in calcrete carbon on land accompanying the transition between glacial and interglacial conditions. This represents a change of about 500–400 GtC Žouter error limits are estimated at 750–200 GtC. . In order to be weathered into dissolved bicarbonate, this would take up an additional 500–400 GtC Ž750–200 GtC. in CO 2 from oceanratmosphere sources. An equivalent amount to the carbonate leaving the caliche reservoir on land may have accumulated in coral reefs and other calcareous marine sediments during the Holocene, liberating an equimolar quantity of CO 2 back into the ocean-atmosphere system as the bicarbonate ion breaks up. q 1999 Elsevier Science B.V. All rights reserved. Keywords: calcrete; carbonate; carbon cycle; LGM; Holocene
1. Introduction Polar ice cores indicate that the levels of greenhouse gases in the Earth’s atmosphere have closely paralleled the changing climate over at least the last 250,000 years Že.g., Barnola et al., 1989.. There is presently widespread interest in understanding the controls on these changes in atmospheric composition. Work has concentrated on producing ocean models to explain the sequestering of larger amounts of carbon into the glacial-age ocean. More recently, )
Corresponding author. E-mail: [email protected].
however, there have been various attempts to provide a more complete picture by estimating how land carbon storage might have changed between glacial and interglacial conditions Že.g., Faure, 1990; Adams et al., 1990; Van Campo et al., 1993; Prentice et al., 1993; Adams and Faure, 1998.. Attention in these papers has concentrated on the amount of organic carbon held in terrestrial vegetation, soils and peats. However, there were other carbon reservoirs and sinks that have not been considered in detail. These include soil carbonate, which often occurs as layers or nodules Žcalcrete. in aridland soils. Various attempts have been made to
0921-8181r99r$ - see front matter q 1999 Elsevier Science B.V. All rights reserved. PII: S 0 9 2 1 - 8 1 8 1 Ž 9 9 . 0 0 0 1 5 - 6
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J.M. Adams, W.M. Post r Global and Planetary Change 20 (1999) 243–256
estimate the size of this reservoir in present-day soils; Schlesinger Ž1982. suggested 780–930 GtC Žs PgC., and Sombroek et al. Ž1993. suggested a figure of 720 GtC. In what may be the most reliable study to date Žincluding an extensive FAO soils database., Batjes Ž1996. and Batjes and Sombroek Ž1997. suggest that the true figure for the uppermost 100 cm of soils is around 700 GtC Žbetween 695 and 748 GtC.. Kern and Schlesinger Ž1992. first explicitly pointed out the long-term potential importance of soil carbonate. Noting that soil carbonate is much more abundant in arid-land soils, and that arid environments were more widespread during the Last Glacial period, Kern and Schlesinger suggest that the global soil carbonate reservoir would have moved in the opposite direction to the organic soil and vegetation carbon reservoir. However, since a loss of soil carbonate involves Žat least temporarily. uptake of CO 2 to dissolve it into bicarbonate, this may have acted as a CO 2 sink in parallel with the increasing land organic carbon record ŽAdams, 1993.. Here we attempt to consider in more detail the possible changes and fate of soil carbonate and the CO 2 that it takes up in weathering, between glacial and interglacial conditions.
2. Method 2.1. Soil carbonate by life zone We used a summary of global soil calcrete carbon content according to Holdridge life zone type, given by Batjes and Sombroek Ž1997. ŽTable 1.. Their table was based upon a large worldwide FAO-ISRIC Ž1989. dataset of soil carbon collected from the literature. In addition to climate and vegetation, soil carbonate content is significantly affected by local hydrology and rock types, but on a global scale such local factors tend to balance out. An uncertainty limit on the mean inorganic carbon content in each life zonerbiome is not given by Batjes and Sombroek, so we are unable to provide rigorous error limits on this component of this calculations. Generally, soil and organic carbon inventories Že.g., Zinke et al., 1984. assume an error limit of around plus or minus 30%, and for purposes of consideration we adopt this
error factor. A 30% error limit also appears to encompass the range of published estimates to date ŽKern and Schlesinger, 1992; Sombroek et al., 1993; Fig. 1.. 2.2. Reconstructing past life zone distributions In order to make estimates of how global soil carbonate carbon would have varied with interglacial cycles, we used the QEN paleovegetation biome figures ŽQEN, 1995; Adams and Faure, 1997.. A conversion scheme showing the lumping of QEN categories into the categories used by Batjes and Sombroek is shown in Table 1. For the presentpotential world, the global total according to the converted scheme is 722 GtC. This agrees well with the figure of ‘around 700 Gt’ found by Batjes Ž1996.. Maps of regional or global vegetation coverage for several time intervals since the Last Glacial Maximum have been produced by several different sources over the past few years Že.g., Crowley, 1995.. The maps presented here ŽFig. 1a–d., while still subject to uncertainties due to sparse data coverage in many areas, represent an approximate ‘consensus’ picture based on published review sources and advice from specialists in Quaternary environments. The sources of evidence used for the maps include: Ž1. plant fossils, Ž2. animal fossils and Ž3. sedimentological indicators. However, there are still significant uncertainties and disagreements among Quaternarists concerning the actual global vegetation cover over the last 20,000 years. The range of scenarios explored below is intended to allow for this element of uncertainty, varying the forest arearforest climate areas by an amount which appears to be plausible on the basis of the available evidence. These different scenarios can be expected to influence overall calculated soil carbonate. 2.3. The steppe-tundra and its soil carbonate content The QEN maps for the Last Glacial MaximumrLate glacial aridity maximum Žapproximately 18,000–14,500 14C ya. show large reductions in forest vegetation, especially in the higher latitudes where the interglacial forests were mostly replaced by ice sheets, polar desert and by the cold-climate steppe tundra vegetation which apparently has no
J.M. Adams, W.M. Post r Global and Planetary Change 20 (1999) 243–256
245
Table 1 Equivalency between QEN ecosystems Žused here. and those of Batjes, and the per-unit area soil carbonate storage in each Ecosystem type
BatjesrHoldridge equivalents
Calcrete, C kgrm2
Ž1. Tropical Rainforest Ž2. Monsoon Forest Ž3. Tropical Woodlands Ž4. Tropical Scrub Ž5. Tropical Semi-Desert Ž6. Tropical Grasslands Ž7. Tropical Desert Ž9. Savanna Ž10. Warm Temperate Forest Ž11. Coastal Giant Conifer Ž12. Tropical Montane Forest Ž13a. Mediterranean Forest Ž13b. Mediterranean Scrub Ž14. Cool Temperate Forest Ž15. South Taiga Ž16. Mid Taiga Ž17. Open Boreal Woodlands Ž18a. Temperate Open Woodland Ž18b. Tempte Scrub Ž19a. MontanerDry Tundra Ž19b. Lowland Tundra Ž20a. Steppic Steppe-Tundra Ž20b. Tundric Steppe-Tundra Ž21. PolarrMountain Desert Ž22. Temperate Desert Ž23. Temperate Semi-Desert Ž24a. Moist Steppe Ž24b. Dry Steppe Ž25. Forest Steppe Ž26. Forest Tundra Ž27. BogrSwamp Ž28. Ice Ž29. Lakerriver
Rain Forest Tropical Seasonal Forest Tropical Dry Forest Tropical Semi-arid Tropical Semi-arid Tropical Semi-arid Hot Desert Tropical Semi-arid Temperate Forest Temperate Forest Temperate Forest Temperate Forest Chapparel Temperate Forest Boreal Forest Boreal Forest Cold parkland Temperate Forest Chapparal Tundra Tundra Cool DesertrSteppe Cool DesertrSteppe Cool Desert Cool Desert Chapparal Steppe Steppe Cold Parkland Cold Parkland not included not included not included
0.5 0.7 3 8 8 8 12 8 1.5 1.5 1.5 3 3 3 2 2 1.5 3 10 1 1 10 10 12.5 12.5 10 10 10 5.5 5.5 0 0 0
exact analogues in the present-day world ŽFrenzel, 1968; Adams and Faure, 1997.. A range of floristic, faunal and pedological evidence shows that the steppe-tundra biome across most of Eurasia was subject to rather arid conditions Žsee Spasskaya, 1992, and discussion by Adams and Faure on the QEN web page., with soil carbonate tending to accumulate much as in a present-day dry steppe or semi-desert environment ŽZelikson, pers. commun., 1995; Morozova et al., 1998.. For example, on the basis of a range of indicators, Morozova et al. Ž1998. conclude that the soils of the eastern European Plain at the Last Glacial Maximum were most similar to Žthough even more arid than. present-day haplic or
arenic calcisols of the Central Asian deserts, with secondary carbonate and gypsum accumulation. Here, we perhaps rather conservatively assign the average of the steppe-tundra to a ‘dry steppe’ category in terms of soil carbonate content, rather than the more carbonate-rich semi-desert. 2.4. Dry and moist ‘limit’ scenarios In many regions, broad-scale paleovegetation reconstructions may be subject to substantial error, due to the sparsely of data, although there is a general consensus among Quaternary vegetation scientists that there was indeed a major Žof the order of 50% or
246 J.M. Adams, W.M. Post r Global and Planetary Change 20 (1999) 243–256 Fig. 1. Ža–d. Reconstructed closed forest and extreme desert conditions Ž‘preferred’ most likely scenario. for the Last Glacial Maximum, early and mid Holocene and present-potential.
J.M. Adams, W.M. Post r Global and Planetary Change 20 (1999) 243–256
Fig. 1 Ž continued ..
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J.M. Adams, W.M. Post r Global and Planetary Change 20 (1999) 243–256
Fig. 1 Ž continued ..
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J.M. Adams, W.M. Post r Global and Planetary Change 20 (1999) 243–256
Fig. 1 Ž continued ..
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J.M. Adams, W.M. Post r Global and Planetary Change 20 (1999) 243–256
more. LGM loss of forest cover from the tropics, together with expansions of desert zones in north Africa, Australia, South America and Eurasia. Although opinions differ on exactly where the desert and forest boundaries would have lain at each timeslice ŽPrentice et al., 1993; Crowley, 1995; Adams, 1995; Webb et al., 1995; Adams and Faure, 1997, 1998., we suggest that our ‘high forest’ and ‘low forest’ scenarios take in most of the range of opinion which is reasonably well supported by the current evidence ŽTable 2.. Note that the emphasis here has been on varying the areas of broad classes according to the Batjes and Sombroek soil calcrete categories
Že.g., temperate forest vs. steppe and non-forest, or desert and semi-desert vs. non-desert. rather than concentrating on the detailed uncertainties for each vegetation category. Our ‘preferred’ scenario ŽTable 2 and Fig. 1a–c. is based on the QEN ecosystem maps of Adams and Faure Ž1997. Žalso available at the QEN web site.. Our ‘high forest’ scenario is more closely in line with the mapping effort of Crowley Ž1995. and the scenario expoused by Colinvaux Ž1987., although we suggest that the evidence for a major reduction in the temperate forest biome Žfor example, van der Hammen and Absy, 1994. is stronger than the evidence for major reduction in the
Table 2 Reconstructed global areas of ecosystem types, from Adams and Faure Ž1997.. Bracketed values are for the low forestrdryer climate and high forestrwetter climate scenarios, respectively. Values are in Mkmy2 Ecosystem code
LGM
8000 years
5000 years
Present pot.
Ž1. TrRfr Ž2. MsnFor Ž3. TrWd Ž4. TrThnSc Ž5. TrSDes Ž6. TrGrs Ž7. TrDes Ž9. Sav Ž10. WarmTemFor Ž11. CtGiantConfr Ž12. TrMontFor Ž13a. MdtnForest Ž13b. MdtnScrub Ž14. CoolTempFor Ž15. SouthTaiga Ž16. Mid Tg Ž17. OpnBWd Ž18a. Temp wdld Ž18b. Temp Scrub Ž19a. Mtnrdry Tu Ž19b. Lowl Tu Ž20a. Steppic St-Tu Ž20b. Tundric St-Tu Ž21. PolarrMtn Des Ž22. TempDes Ž23. TempSDes Ž24a. Moist Steppe Ž24b. Dry Steppe Ž25. ForSt Ž26. ForTu Ž27. BogrSwamp Ž28. Ice Ž29. Lakerriver Global area
5.56 1.9 1.19 3.33 4.64 10.50 13.83 12.25 0.76 0 0.54 0 0.17 0.2 0 0.86 2.50 4.35 2.91 0.73 0.13 11.82 3.34 15.52 12.50 7.10 0.50 6.50 1.3 – – 36.09 1.3 161
Ž4–8. 16.38 Ž1.2–3.5. 4.5 Ž1.19–2.69. 6.62 Ž5.33–4.33. 5.99 Ž4.64–3.15. 9.43 Ž12–10.50. 7.64 Ž13.83–11.83. 0.39 Ž9.25–11.83. 11.00 Ž0.2–1.0. 4.55 0.35 Ž0.54–0.74. 1.20 0.68 Ž0.17–0.27. 0.66 Ž0.1–1.5. 9.75 Ž0–0.5. 6.53 Ž0.86–2.5. 7.17 Ž2.5–2.5. 4.84 Ž4.35–4.35. 2.06 Ž2.91–2.91. 0.46 Ž0.73–0.73. 2.57 Ž0.13–1.9. 1.80 Ž11.82–11.32. 0 Ž3.34–4.32. 0 Ž15.52–8.5. 1.95 Ž12.5–8.5. 0.20 Ž7.10–6.60. 8.16 Ž0.2–1.0. 3.50 Ž7.5–5.50. 3.93 Ž0.7–2.5. 1.20 2.71 1.82 19.86 2.4 150
Ž13.38–18.00. 15.77 Ž3.00–5.00. 4.5 Ž5.00–6.62. 6.17 Ž8.99–5.49. 7.49 Ž9.43–8.43. 10.97 Ž7.64–7.64. 3.89 Ž0.39–0.39. 1.69 Ž11.00–11.50. 11.82 Ž4.0–5.09. 4.59 Ž0.25–0.35. 0.38 Ž1.20–1.20. 1.10 Ž0.50–0.68. 0.78 Ž0.56–0.66. 0.20 Ž9.20–11.00. 11.27 Ž5.53–7.00. 8.87 Ž6.17–7.60. 5.68 Ž4.84–6.53. 3.83 Ž2.06–2.06. 2.26 Ž0.46–0.46. 0.46 Ž2.57–2.57. 3.01 Ž1.00–2.00. 2.67 0 0 Ž1.95–1.75. 2.22 Ž0.50–0.10. 0.20 Ž6.01–5.61. 8.70 Ž3.00–4.50. 4.71 Ž4.93–2.93. 3.03 Ž1.20–1.20. 1.57 Ž1.50–3.00. 1.08 Ž1.82–1.82. 2.30 15.91 2.7 149
Ž13.0–16.5. 12.47 Ž3.00–5.00. 2.96 Ž4.50–6.17. 6.17 Ž9.49–6.99. 10.85 Ž10.97–9.97. 6.98 Ž4.89–3.89. 2.79 Ž0.39–2.00. 11.73 Ž11.82–10.82. 8.45 Ž4.0–5.09. 3.45 Ž0.35–0.38. 0.38 Ž1.10–1.10. 0.96 Ž0.50–0.78. 0.41 Ž0.40–0.20. 0.2 Ž11.00–12.5. 10.82 Ž7.87–9.00. 5.7 Ž6.68–5.68. 6.26 Ž3.83–4.83. 3.22 Ž1.26–2.26. 2.03 Ž0.46–1.46. 1.18 Ž3.01–3.33. 3.33 Ž2.67–2.67. 4.49 0 0 Ž2.22–2.00. 2.08 Ž0.50–0.10. 2.42 Ž7.20–6.80. 7.72 Ž4.81–5.71. 4.87 Ž6.30–4.03. 3.39 Ž1.57–1.57. 1.96 Ž1.08–1.08. – Ž2.35–2.35. 2.43 15.91 2.65 149.85
J.M. Adams, W.M. Post r Global and Planetary Change 20 (1999) 243–256
tropical forest biome, due to the relative abundance of a range of forms of palaeoevidence from the temperate regions. For our ‘low forest’ scenario we suggest even greater reductions in forest area than is considered most likely by Adams and Faure, and more in line with the views of authors such as Frenzel Ž1968. and Clapperton Ž1993..
3. Results and discussion If for consideration one simply compares the range of possible values ŽFig. 2. which would result from a "30% error limit on the average soil carbonate content in each biome, the glacial-to-Holocene change could in principle involve no change or even a slight increase in soil carbonate, or as much as a 1000 Gt decrease. However, if due to sampling error a particular biome in the present-day world has Žfor example. a 30% lower soil carbonate content on average than the preferred figure which Batjes and Sombroek Ž1997. give, this error will presumably follow in the same direction for both glacial and interglacial conditions, rather than making it possible to compare an extreme q30% error for one timeslice against a y30% error for another timeslice. This constraint will tend to ‘pull in’ the true limits of uncertainty on the glacial-to-Holocene change in soil carbonate carbon. At a more general level, the width of error limits should not be allowed to distract from the fact that a several hundred Gt decrease in carbonate carbon remains much more likely from the general biome distributions, and from what is known of their soil carbonate content from present-day nearest analogs. A comparison of the LGM timeslice with the three Holocene timeslices ŽTable 3, Fig. 2. suggests that a major Žaround 40–50%. decrease in soil carbonate carbon occurred between LGM and early-tomid Holocene conditions, due to the greatly decreased areas of arid and semi-arid biomesrclimate zones ŽTable 3.. For example, if one takes the ‘preferred’ figure for the LGM, and compares it to the preferred figures for the Holocene ŽFig. 2., the shift is towards about 400–500 Gt less carbon. As an uppermost estimate for the shift, subtracting the ‘low forestrdry world’ q30% error scenario for the LGM Ž1497 GtC. from a ‘high forest, wet world’ q30%
251
error scenario for the 8000 14C ya, scenario Ž731 GtC. gives a larger shift of 766 GtC. As a lowermost estimate, a shift from the y30% error ‘high forestrwet world’ scenario for the LGM Ž697 GtC. to the ‘low forestrdry world’ y30% error scenario for the early Holocene Ž468 GtC., would give a relatively low shift of 229 GtC. Thus, the overall possible range for the ‘true’ figure of the inorganic carbon storage change can be seen as lying somewhere between 766 and 229 Gt, with the figure most likely to have been around 400–500 Gt considering the per-unit-area estimates given by Batjes and Sombroek, and the approximate consensus global biome distributions produced by QEN ŽFig. 1a–d.. This shift to less carbon in soil carbonates with the onset of interglacial conditions would have been associated with a major efflux of soil carbonate, most of it presumably chemically weathered in the following reaction, partly catalysed by biological activity in the soil; calcium carbonateq carbon dioxideq water | dissolved calcium ions q dissolved bicarbonate ions CaCO 3 q CO 2 q H 2 O | Ca2qq 2 Ž HCO 3 .
y
Smaller amounts of magnesium carbonate present in soil carbonates would have been dissolved in the same manner, and a minor fraction Žjudging from the ratio of sulphate to carbonate in present-day groundwater and river water compositions; Probst, 1994. weathered by sulphuric acids to give dissolved sulphate salts. This latter reaction would release a small proportion of CO 2 to the atmosphere. In the present-day world, weathered bicarbonate enters groundwater or is washed down in surface flow to rivers, and ultimately the oceans ŽProbst, 1994.. In order to be weathered into dissolved bicarbonate, the diminishing reservoir of soil carbonate would have to take up an additional 400–500 Žouter limits about 750–200. GtC as CO 2 derived from oceanratmosphere sources. If 2.1 GtC equals approximately 1 ppm CO 2 , this represents the equivalent of 190–238 ppm of CO 2 in the atmosphere. The soil organic carbon reservoir under LGM and late glacial conditions also seems to have been much less
252
J.M. Adams, W.M. Post r Global and Planetary Change 20 (1999) 243–256
Fig. 2. Calculated global GtC in soil carbonates. Upper and lower limits are shown for ‘low forest’ and ‘high forest’ vegetation distributions, respectively. Outermost limits allow for uncertainty in mean soil carbonate content in each life zone. Previously published present-day estimates by Schlesinger and by Batjes are also shown for comparison.
than during the Holocene and would thus have been unable to supply this extra quantity of carbon by oxidation ŽAdams and Faure, 1998.. As Batjes and Sombroek Ž1997. point out, the detailed dynamics of soil carbonate accumulation and loss during a future switchover in climate and water regime are highly uncertain. The same can be said of past changes, and our calculations here as-
sume that a rough equilibrium state had been achieved given the climate and vegetation conditions at each timeslice—an assumption which may be difficult to test given the present state of knowledge of soil carbonate dynamics, but which might not be too inaccurate considering the thousands of years for which the general climatic and vegetation conditions at each timeslice persisted.
J.M. Adams, W.M. Post r Global and Planetary Change 20 (1999) 243–256
253
Table 3 Calculated global soil carbon storage in carbonate. Values are in t = 10 8. Preferred scenario is outside the parentheses, while the first value in parentheses is for the ‘wetter climate’ scenario with more forest, the second is for the ‘drier climate’ scenario with less forest and more arid and semi-arid vegetation Ecosystem code
LGM
8000 years
5000 years
Present pot.
Ž1. TrRfr Ž2. MsnFor Ž3. TrWd Ž4. TrThnSc Ž5. TrSDes Ž6. TrGrs Ž7. TrDes Ž9. Sav Ž10. WarmTemFor Ž11. CtGiantConfr Ž12. TrMontFor Ž13a. MdtnForest Ž13b. MdtnScrub Ž14. CoolTempFor Ž15. SouthTaiga Ž16. Mid Tg Ž17. OpnBWd Ž18a. Temp wdld Ž18b. Temp Scrub Ž19a. Mtnrdry Tu Ž19b. Lowl Tu Ž20a. Steppic St-Tu Ž20b. Tundric St-Tu Ž21. PolarrMontane Desert Ž22. Temperate Desert Ž23. Temperate Semi-Desert Ž24a. Moist Steppe Ž24b. Dry Steppe Ž25. Forest Steppe Ž26. Forest Tundra Ž27. BogrSwamp Ž28. Ice Ž29. Lakerriver Total carbon
27.8 Ž37.5–20. 13.5 Ž24.5–8.4. 35.7 Ž80.7–35.7. 266.4 Ž346.4–426.4. 371.2 Ž252–371.2. 840 Ž840–960. 1659.6 Ž1419–1740. 980 Ž946–740. 11.4 Ž11.4–740. 0 8.1 Ž11.1–7.5. 0 5.1 Ž8.1–5.1. 6 Ž45–3. 0 24 Ž50–17.2. 37.5 Ž52.5–22.5. 130.5 Ž130.5–130.5. 291 Ž291–291. 7.3 Ž7.3–7.3. 1.3 Ž19–1.3. 1182 Ž1182–1182. 334 Ž432–334. 1940 Ž1190–1940. 1562.5 Ž1062–1562. 710 Ž810–910. 50 Ž100–20. 650 Ž550–750. 71.5 Ž71.5–38.5. 0 0 0 0 11,216
81.9 Ž90–66.9. 31.5 Ž35–21. 198.1 Ž198.5–150. 479 Ž439–639.2. 754 Ž674–834.4. 611.2 Ž611–611.2. 46.8 Ž14.4–46.4. 869 Ž920–880. 68.2 Ž68.2–60. 5.25 Ž5.2–3.75. 18 Ž22.5–18. 20.4 Ž23.4–15. 19.8 Ž16.8–16.8. 292.5 Ž330–276. 130.6 Ž140–110.6. 143.4 Ž152–123.4. 72.6 Ž78.4–87.6. 61.8 Ž61.8–91.8. 46 Ž46–446. 25.7 Ž15.7–25.7. 18 Ž25–10. 0 0 243.7 Ž218.7–243.7. 25 Ž12.5–62.5. 816 Ž461–601. 350 Ž400–300. 393 Ž293–693. 66 Ž121–121. 149.0 Ž165–137.5. 0 0 0 6,038
78.8 Ž81–65. 31.5 Ž38.5–22.4. 185.6 Ž210.3–135.2. 479.2 Ž559.5–839.6. 754.4 Ž717.6–877.6. 611.2 Ž311.2–471.2. 46.8 Ž46.8–46.8. 865.6 Ž705.6–945.6. 68.25 Ž76.3–60. 5.7 Ž5.7–5.25. 18 Ž37.5–30. 23.4 Ž23.4–15. 6 Ž6–12. 338.1 Ž375–330. 177.6 Ž180–157.4. 113.6 Ž100–133.6. 60.15 Ž60–57.4. 67.8 Ž67.8–37.8. 46 Ž50–246. 33.1 Ž33.3–33. 26.7 Ž26.7–21. 0 0 277.6 Ž250–277.6. 25 Ž12.5–62.5. 870 Ž600–820. 471 Ž671–471. 303 Ž603–550. 86.3 Ž86.3–86.35. 59.4 Ž86.9–49.5. 0 0 0 6,086
62.3 20.7 185.1 868.2 558.6 223.2 1407 676.6 51.8 5.7 14.5 12.4 6 324.1 114.4 125.6 48.1 60.8 118 33.3 44.9 0 0 260 302.5 772 487 339 107.8 0 0 0 0 7,229
3.1. Other interacting carbonate sinks r sources It is unclear how much of the ‘missing’ 200–750 Gt carbon taken up by postglacial soil carbonate weathering is still held in the oceans or in groundwater, and what proportion of it cycled rapidly through the system, depositing carbonate on sea floors or coral reefs and re-entering the atmosphere. The diminution of the soil carbonate reservoir adds to the problems involved with modelling the glacial carbon cycle, which already cannot account for the source of CO 2 needed to replenish organic soil and vegetation
stocks on land during the interglacial ŽAdams and Faure, 1998.. One possibility is that the diminishing caliche reservoir on land has been approximately balanced by the accumulating coral reef reservoir during the Holocene: an estimate by Faure et al. Ž1988. suggests that about 500 Gt inorganic C has accumulated in coral reefs since the start of the Holocene Žallowing for 100 Gt internally recycled carbonate due to bioerosion; H. Faure, pers. commun., 1998.. Since coral reefs grow by splitting dissolved bicarbonate to release insoluble calcium carbonate and CO 2 Žused
254
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in photosynthesis by coralline algae., this mechanism could approximately balance the ‘extra’ carbon taken up in caliche weathering on land. The CO 2 released may partly have been retained as organic matter within the buried coral, but after the death of the coral organism most of it would presumably have been released as CO 2 into the general ocean–atmosphere system Žgiven the fairly low organic matter content of coral rock.. Hence, overall, there may be little or no net flux of atmospheric CO 2 into the ocean’s dissolved carbonate reservoir, with calcretederived bicarbonate merely acting as a temporary ‘ vehicle’ for CO 2 . In addition to corals, there is the quantity of carbonates precipitating in muds and oozes on the continental shelves Žwhich increased in area as the sea level rose during deglaciation. and the continental slopes, although no specific estimate of this accumulation since the beginning of the Holocene is available. However, during the early Holocene, the rapid initial loss of calcrete might not have been fully counterbalanced by the growth rate of reefs, and the slowness of sea level rise may have delayed the main episodes of abiotic shelf carbonate precipitation until well into the Holocene. Other subaerial bicarbonate sources need to be considered in addition to soil carbonates, to place our estimates in their true context. The other sources seem likely to have changed in the same direction as the calcretes, taking up more CO 2 and sending more bicarbonate to the oceans under interglacial conditions. It is probable that the quantity of carbon taken up each year by silicate weathering Žalso ultimately the source of soil carbonate. was greatly affected by the glacial-to-interglacial climate and vegetation change ŽAdams, 1993, 1995; Adams and Faure, 1997. perhaps adding several further hundreds of gigatonnes to the cumulative extra amount of CO 2 taken up by the land surface, as opposed to any corresponding time interval of the last glacial Oxygen Isotope Stage 2. This would further have added to the burden of dissolved carbonate entering the oceans, whose fate needs to be explained. Karst areas of limestone on land weather rapidly and provide a further major sink of CO 2 and source of dissolved carbonate to the oceans, which would have varied between glacial and interglacial conditions with the large changes in moisture conditions.
Present-day uptake of carbon by karst carbonate weathering may be of the order of 0.2–0.6 GtCryr ŽIGBP, 1998.. As Ludwig et al. Ž1998. have suggested, carbonate weathering influx Žand rates of CO 2 uptake as bicarbonate leaving for the ocean. on dead reefs on exposed continental shelves during the Last Glacial might have partly offset the change resulting from the switch from dry to moist climates around the world. However, it is necessary to bear in mind that sea levels took much longer to rise than climate took to change; large areas of coral reef would still have been exposed to subaerial weathering under the renewed warmer and moister climates of the interglacial. This itself should have provided a great increase in flux of carbonate to the oceans between glacial and interglacial conditions; an additional flux which has not been considered in this paper and which awaits further work. 3.2. The context of the atmospheric CO2 and ocean carbonate record Interpretation of the ocean carbonate Žlysocline. record ŽBroecker and Sanyal, 1997., needs to take into account the range of different processes including calcrete weathering, which may have produced a rapid increase in efflux Žof the order of hundreds of gigatonnes. of carbon in dissolved bicarbonate to the oceans during the several thousand year period of deglaciation. The role of groundwater accumulation of bicarbonate and re-deposition of carbonate deeper within the soil system needs to be considered, but it is unlikely that this can account for more than a small fraction of the loss of carbonate from calcretes. The general record of the balance between carbonate preservation and dissolution in ocean sediments suggests that during deglaciation there was an increase in carbonate preservation at greater depths Ža deepening of the lysocline, the level below which carbonate redissolves.. This is consistent with withdrawal of CO 2 from the ocean by a regenerating land biota ŽBroecker and Sanyal, 1997., but it may also partly be a product of more carbonate from weathering being ‘shuttled’ to the oceans as bicarbonate and released by carbonate precipitation, with the CO 2 from this dissociation ultimately being absorbed as ecosystem organic carbon storage back on
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land. It is important to bear in mind, however, that variations in ocean primary productivity and in watermass circulation patterns could also be important in controlling changes in the lysocline. Furthermore, the resulting flux—involving hundreds of Gt extra CO 2 transport by bicarbonate during a several thousand year period—needs to be incorporated into carbon cycle models. A decrease of 750–200 GtC in soil calcrete would require some 750–200 GtC of CO 2 to dissolve it into bicarbonate. This soil calcrete flux alone is equivalent to between 357 and 95 ppm atmospheric CO 2 Žif one assumes a relationship of 2.1 GtC s 1 ppm CO 2 .. This is enough to have had a significant effect on the postglacial accumulation of CO 2 in the atmosphere if there was a delay in the recycling of this CO 2 back out of the oceans taking more than a few hundred years. To this must be added the effect of a probable major increase in subaerial weathering of exposed reefs, older carbonate rocks and silicate rocks as climates around the world switched to their interglacial state.
Acknowledgements The advice of Hugues Faure, Ellen Thomas, Lloyd Keigwin and our two anonymous reviewers is much appreciated. Jonathan Adams was supported by the Lockheed-Martin on a Eugene P. Wigner Fellowship Program at Oak Ridge National Laboratory. Research sponsored by the U.S. Department of Energy, Carbon Dioxide Research Program, Environmental Sciences Division, Office of Biological and Environmental Research, under contract DE-AC05-96OR22464 with Lockheed Martin Energy Research Corp.
References Adams, J.M., 1993. Caliche carbonate and the carbon cycle. Nature 361, 213–214. Adams, J.M., 1995. The role of terrestrial ecosystems in glacial to interglacial changes in the global carbon cycle. PhD thesis. Universite´ D’Aix-Marseille II, France. Adams, J.M., Faure, H., 1997. Palaeovegetation maps of the Earth during the Last Glacial Maximum, and the early and mid
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Holocene: an aid to archaeologists. J. Archaeol. Sci. 24, 623–647. Adams, J.M., Faure, H., 1998. A new estimate of changing carbon storage on land since the last glacial maximum. Global Planet. Change 16, 3–24. Adams, J.M., Faure, H., Faure-Denard, L., McGlade, J.M., Woodward, F.I., 1990. Increases in terrestrial carbon storage from the Last Glacial Maximum to the present. Nature 348, 711– 714. Barnola, J.M., Raynaud, D., Koroketvich, Y.S., Lorius, C., 1989. Vostok ice core: a 160,000 year record of atmospheric CO 2 . Nature 329, 408–414. Batjes, N.H., 1996. Total carbon and nitrogen in soils of the world. Eur. J. Soil Sci. 47, 151–163. Batjes, N.H., Sombroek, W.G., 1997. Possibilities for carbon sequestration in tropical and subtropical soils. Global Change Biol. 3, 161–173. Broecker, W.S., Sanyal, A., 1997. Magnitude of the CaCO 3 dissolution events marking the onset of times of glaciation. Paleoceanography 12, 530–532. Clapperton, C.M., 1993. Quaternary Geology and Geomorphology of South America. Elsevier, Amsterdam. Colinvaux, P.A., 1987. Environmental history of the Amazon Basin, Quaternary of South America and Antarctic Peninsula 5 A.A. Balkema, Rotterdam, Brookfield. Crowley, T.J., 1995. Ice age terrestrial carbon changes revisited. Global Biogeochem. Cycles, in press. FAO-ISRIC Soil Database ŽSDB., 1989. World Soil Resources Report 64. Food and Agriculture Organization, Rome. Faure, H., 1990. Changes in the continental reservoir of carbon. Palaeogeogr. Palaeoclim. Palaeoecol. 82, 47–52. Faure, H., Giresse, P., Pirazzoli, P.A., 1988. Bilan global des accumulations organogenes post-glaciaires. Colloque de Biosedimentologie, Volume des Resumes A.P.F., A.S.F., S.G.F., Marseille, October, 1988. Frenzel, B., 1968. The Pleistocene vegetation of northern Eurasia. Science 161, 637–639. IGBP, 1998. Project 379. Karst processes and the carbon cycle. In: Daoxin, Y., Zaihua, L. ŽEds.., Annual Report. Kern, R., Schlesinger, W.H., 1992. Carbon stores on vegetation. Nature 357, 447–449. Ludwig, W., Amiotte-Suchet, P., Munhoven, G., Probst, J.L., 1998. Atmospheric CO 2 consumption by continental erosion: present-day controls and implications for the Last Glacial Maximum. Global Planet. Change 16, 34–60. Morozova, T.D., Velichko, A.A., Dlussky, K.G., 1998. Organic carbon content in the late Pleistocene and Holocene fossil soils Žreconstruction for eastern Europe.. Global Planet. Change 16, 131–151. Prentice, I.C., Sykes, M.T., Lautenschlager, M., Harrison, S.P., Denissenko, O., Bartlein, P., 1993. Modelling global vegetation patterns and terrestrial carbon storage at the last glacial maximum. Global Ecol. Biogeogr. Lett. 3, 67–76. Probst, J-L., 1994. Geochimie et Hydrologie de L’Erosion Continentale. Science Geologiiques. Memoires 94 ŽDoctoral Thesis.. University of Strasbourg. QEN, 1995. The Quaternary Environments Network Atlas and
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Review of Palaeovegetation during the last 20,000 years. World Wide Web address: http:´www.esd.ornl.gov.rprojectsr genradams1.html Schlesinger, W.H., 1982. Carbon storage in the caliche of the arid world: a case study from Arizona. Soil Sci. 133, 247–255. Sombroek, W.G., Nachtergaele, F.O., Hebel, A., 1993. Amounts, dynamics and sequestrations of carbon in tropical and subtropical soils. Ambio 22, 417–426. Spasskaya, I.I., 1992. Dominant geomorphic processes during the maximum cooling of the last glaciation. In: Frenzel, B., Pecsi, B., Velichko, A.A. ŽEds.., Atlas of Palaeoclimates and Palaeoenvironments of the Northern Hemisphere. INQUAr Hungarian Academy of Sciences, Budapest.
Van Campo, E., Guiot, J., Peng, C., 1993. A data-based re-appraisal of the terrestrial carbon budget at the last glacial maximum. Global Planet. Change 8, 189–201. van der Hammen, T., Absy, M.L., 1994. Amazonia during the last Glacial. Palaeogeogr. Palaeoclim. Palaeoecol. 109, 247–261. Webb, R., et al., 1995. Bibliography and inventory of sites with information for constructing digital global maps of vegetation distribution 18,000 yr. B.P., NOAA Palaeoclimatol. Publ. Ser. Rep. Nat. Geophys. Data Cent., Boulder, CO, in preparation. Zinke, P.J., Stangenberger, A.G., Post, W.M., Emanuel, W.R., Olson, J.S., 1984. Worldwide organic soil carbon and nitrogen data. Environmental Sciences Division, publication no. 2212. Oak Ridge National LaboratoryrUS Department of Energy.
A preliminary estimate of changing calcrete carbon storage on land since the Last Glacial Maximum J.M. Adams b
a,b,)
, W.M. Post
a
a MS 6335, EnÕironmental Sciences DiÕision, Building 1509, Oak Ridge National Laboratory, Oak Ridge, TN 37831, USA Department of EnÕironmental and Geographical Sciences, UniÕersity of Adelaide, Napier Building, Adelaide, SA 5005, Australia
Abstract The glacial-to-interglacial shift in land carbon storage is important in understanding the global carbon cycle and history of the climate system. While organic carbon storage on land appears to have been much less than present during the cold, dry glacial maximum, calcrete Žsoil carbonate. carbon storage would have been greater. Here we attempt a global estimation of this change; we use published figures for present soil carbonate by biome to estimate changing global soil carbonate storage, on the basis of reconstruction of vegetation areas for four timeslices since the Last Glacial Maximum. It appears that there would most likely have been around a 30–45% decrease in calcrete carbon on land accompanying the transition between glacial and interglacial conditions. This represents a change of about 500–400 GtC Žouter error limits are estimated at 750–200 GtC. . In order to be weathered into dissolved bicarbonate, this would take up an additional 500–400 GtC Ž750–200 GtC. in CO 2 from oceanratmosphere sources. An equivalent amount to the carbonate leaving the caliche reservoir on land may have accumulated in coral reefs and other calcareous marine sediments during the Holocene, liberating an equimolar quantity of CO 2 back into the ocean-atmosphere system as the bicarbonate ion breaks up. q 1999 Elsevier Science B.V. All rights reserved. Keywords: calcrete; carbonate; carbon cycle; LGM; Holocene
1. Introduction Polar ice cores indicate that the levels of greenhouse gases in the Earth’s atmosphere have closely paralleled the changing climate over at least the last 250,000 years Že.g., Barnola et al., 1989.. There is presently widespread interest in understanding the controls on these changes in atmospheric composition. Work has concentrated on producing ocean models to explain the sequestering of larger amounts of carbon into the glacial-age ocean. More recently, )
Corresponding author. E-mail: [email protected].
however, there have been various attempts to provide a more complete picture by estimating how land carbon storage might have changed between glacial and interglacial conditions Že.g., Faure, 1990; Adams et al., 1990; Van Campo et al., 1993; Prentice et al., 1993; Adams and Faure, 1998.. Attention in these papers has concentrated on the amount of organic carbon held in terrestrial vegetation, soils and peats. However, there were other carbon reservoirs and sinks that have not been considered in detail. These include soil carbonate, which often occurs as layers or nodules Žcalcrete. in aridland soils. Various attempts have been made to
0921-8181r99r$ - see front matter q 1999 Elsevier Science B.V. All rights reserved. PII: S 0 9 2 1 - 8 1 8 1 Ž 9 9 . 0 0 0 1 5 - 6
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estimate the size of this reservoir in present-day soils; Schlesinger Ž1982. suggested 780–930 GtC Žs PgC., and Sombroek et al. Ž1993. suggested a figure of 720 GtC. In what may be the most reliable study to date Žincluding an extensive FAO soils database., Batjes Ž1996. and Batjes and Sombroek Ž1997. suggest that the true figure for the uppermost 100 cm of soils is around 700 GtC Žbetween 695 and 748 GtC.. Kern and Schlesinger Ž1992. first explicitly pointed out the long-term potential importance of soil carbonate. Noting that soil carbonate is much more abundant in arid-land soils, and that arid environments were more widespread during the Last Glacial period, Kern and Schlesinger suggest that the global soil carbonate reservoir would have moved in the opposite direction to the organic soil and vegetation carbon reservoir. However, since a loss of soil carbonate involves Žat least temporarily. uptake of CO 2 to dissolve it into bicarbonate, this may have acted as a CO 2 sink in parallel with the increasing land organic carbon record ŽAdams, 1993.. Here we attempt to consider in more detail the possible changes and fate of soil carbonate and the CO 2 that it takes up in weathering, between glacial and interglacial conditions.
2. Method 2.1. Soil carbonate by life zone We used a summary of global soil calcrete carbon content according to Holdridge life zone type, given by Batjes and Sombroek Ž1997. ŽTable 1.. Their table was based upon a large worldwide FAO-ISRIC Ž1989. dataset of soil carbon collected from the literature. In addition to climate and vegetation, soil carbonate content is significantly affected by local hydrology and rock types, but on a global scale such local factors tend to balance out. An uncertainty limit on the mean inorganic carbon content in each life zonerbiome is not given by Batjes and Sombroek, so we are unable to provide rigorous error limits on this component of this calculations. Generally, soil and organic carbon inventories Že.g., Zinke et al., 1984. assume an error limit of around plus or minus 30%, and for purposes of consideration we adopt this
error factor. A 30% error limit also appears to encompass the range of published estimates to date ŽKern and Schlesinger, 1992; Sombroek et al., 1993; Fig. 1.. 2.2. Reconstructing past life zone distributions In order to make estimates of how global soil carbonate carbon would have varied with interglacial cycles, we used the QEN paleovegetation biome figures ŽQEN, 1995; Adams and Faure, 1997.. A conversion scheme showing the lumping of QEN categories into the categories used by Batjes and Sombroek is shown in Table 1. For the presentpotential world, the global total according to the converted scheme is 722 GtC. This agrees well with the figure of ‘around 700 Gt’ found by Batjes Ž1996.. Maps of regional or global vegetation coverage for several time intervals since the Last Glacial Maximum have been produced by several different sources over the past few years Že.g., Crowley, 1995.. The maps presented here ŽFig. 1a–d., while still subject to uncertainties due to sparse data coverage in many areas, represent an approximate ‘consensus’ picture based on published review sources and advice from specialists in Quaternary environments. The sources of evidence used for the maps include: Ž1. plant fossils, Ž2. animal fossils and Ž3. sedimentological indicators. However, there are still significant uncertainties and disagreements among Quaternarists concerning the actual global vegetation cover over the last 20,000 years. The range of scenarios explored below is intended to allow for this element of uncertainty, varying the forest arearforest climate areas by an amount which appears to be plausible on the basis of the available evidence. These different scenarios can be expected to influence overall calculated soil carbonate. 2.3. The steppe-tundra and its soil carbonate content The QEN maps for the Last Glacial MaximumrLate glacial aridity maximum Žapproximately 18,000–14,500 14C ya. show large reductions in forest vegetation, especially in the higher latitudes where the interglacial forests were mostly replaced by ice sheets, polar desert and by the cold-climate steppe tundra vegetation which apparently has no
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245
Table 1 Equivalency between QEN ecosystems Žused here. and those of Batjes, and the per-unit area soil carbonate storage in each Ecosystem type
BatjesrHoldridge equivalents
Calcrete, C kgrm2
Ž1. Tropical Rainforest Ž2. Monsoon Forest Ž3. Tropical Woodlands Ž4. Tropical Scrub Ž5. Tropical Semi-Desert Ž6. Tropical Grasslands Ž7. Tropical Desert Ž9. Savanna Ž10. Warm Temperate Forest Ž11. Coastal Giant Conifer Ž12. Tropical Montane Forest Ž13a. Mediterranean Forest Ž13b. Mediterranean Scrub Ž14. Cool Temperate Forest Ž15. South Taiga Ž16. Mid Taiga Ž17. Open Boreal Woodlands Ž18a. Temperate Open Woodland Ž18b. Tempte Scrub Ž19a. MontanerDry Tundra Ž19b. Lowland Tundra Ž20a. Steppic Steppe-Tundra Ž20b. Tundric Steppe-Tundra Ž21. PolarrMountain Desert Ž22. Temperate Desert Ž23. Temperate Semi-Desert Ž24a. Moist Steppe Ž24b. Dry Steppe Ž25. Forest Steppe Ž26. Forest Tundra Ž27. BogrSwamp Ž28. Ice Ž29. Lakerriver
Rain Forest Tropical Seasonal Forest Tropical Dry Forest Tropical Semi-arid Tropical Semi-arid Tropical Semi-arid Hot Desert Tropical Semi-arid Temperate Forest Temperate Forest Temperate Forest Temperate Forest Chapparel Temperate Forest Boreal Forest Boreal Forest Cold parkland Temperate Forest Chapparal Tundra Tundra Cool DesertrSteppe Cool DesertrSteppe Cool Desert Cool Desert Chapparal Steppe Steppe Cold Parkland Cold Parkland not included not included not included
0.5 0.7 3 8 8 8 12 8 1.5 1.5 1.5 3 3 3 2 2 1.5 3 10 1 1 10 10 12.5 12.5 10 10 10 5.5 5.5 0 0 0
exact analogues in the present-day world ŽFrenzel, 1968; Adams and Faure, 1997.. A range of floristic, faunal and pedological evidence shows that the steppe-tundra biome across most of Eurasia was subject to rather arid conditions Žsee Spasskaya, 1992, and discussion by Adams and Faure on the QEN web page., with soil carbonate tending to accumulate much as in a present-day dry steppe or semi-desert environment ŽZelikson, pers. commun., 1995; Morozova et al., 1998.. For example, on the basis of a range of indicators, Morozova et al. Ž1998. conclude that the soils of the eastern European Plain at the Last Glacial Maximum were most similar to Žthough even more arid than. present-day haplic or
arenic calcisols of the Central Asian deserts, with secondary carbonate and gypsum accumulation. Here, we perhaps rather conservatively assign the average of the steppe-tundra to a ‘dry steppe’ category in terms of soil carbonate content, rather than the more carbonate-rich semi-desert. 2.4. Dry and moist ‘limit’ scenarios In many regions, broad-scale paleovegetation reconstructions may be subject to substantial error, due to the sparsely of data, although there is a general consensus among Quaternary vegetation scientists that there was indeed a major Žof the order of 50% or
246 J.M. Adams, W.M. Post r Global and Planetary Change 20 (1999) 243–256 Fig. 1. Ža–d. Reconstructed closed forest and extreme desert conditions Ž‘preferred’ most likely scenario. for the Last Glacial Maximum, early and mid Holocene and present-potential.
J.M. Adams, W.M. Post r Global and Planetary Change 20 (1999) 243–256
Fig. 1 Ž continued ..
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Fig. 1 Ž continued ..
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J.M. Adams, W.M. Post r Global and Planetary Change 20 (1999) 243–256
Fig. 1 Ž continued ..
249
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more. LGM loss of forest cover from the tropics, together with expansions of desert zones in north Africa, Australia, South America and Eurasia. Although opinions differ on exactly where the desert and forest boundaries would have lain at each timeslice ŽPrentice et al., 1993; Crowley, 1995; Adams, 1995; Webb et al., 1995; Adams and Faure, 1997, 1998., we suggest that our ‘high forest’ and ‘low forest’ scenarios take in most of the range of opinion which is reasonably well supported by the current evidence ŽTable 2.. Note that the emphasis here has been on varying the areas of broad classes according to the Batjes and Sombroek soil calcrete categories
Že.g., temperate forest vs. steppe and non-forest, or desert and semi-desert vs. non-desert. rather than concentrating on the detailed uncertainties for each vegetation category. Our ‘preferred’ scenario ŽTable 2 and Fig. 1a–c. is based on the QEN ecosystem maps of Adams and Faure Ž1997. Žalso available at the QEN web site.. Our ‘high forest’ scenario is more closely in line with the mapping effort of Crowley Ž1995. and the scenario expoused by Colinvaux Ž1987., although we suggest that the evidence for a major reduction in the temperate forest biome Žfor example, van der Hammen and Absy, 1994. is stronger than the evidence for major reduction in the
Table 2 Reconstructed global areas of ecosystem types, from Adams and Faure Ž1997.. Bracketed values are for the low forestrdryer climate and high forestrwetter climate scenarios, respectively. Values are in Mkmy2 Ecosystem code
LGM
8000 years
5000 years
Present pot.
Ž1. TrRfr Ž2. MsnFor Ž3. TrWd Ž4. TrThnSc Ž5. TrSDes Ž6. TrGrs Ž7. TrDes Ž9. Sav Ž10. WarmTemFor Ž11. CtGiantConfr Ž12. TrMontFor Ž13a. MdtnForest Ž13b. MdtnScrub Ž14. CoolTempFor Ž15. SouthTaiga Ž16. Mid Tg Ž17. OpnBWd Ž18a. Temp wdld Ž18b. Temp Scrub Ž19a. Mtnrdry Tu Ž19b. Lowl Tu Ž20a. Steppic St-Tu Ž20b. Tundric St-Tu Ž21. PolarrMtn Des Ž22. TempDes Ž23. TempSDes Ž24a. Moist Steppe Ž24b. Dry Steppe Ž25. ForSt Ž26. ForTu Ž27. BogrSwamp Ž28. Ice Ž29. Lakerriver Global area
5.56 1.9 1.19 3.33 4.64 10.50 13.83 12.25 0.76 0 0.54 0 0.17 0.2 0 0.86 2.50 4.35 2.91 0.73 0.13 11.82 3.34 15.52 12.50 7.10 0.50 6.50 1.3 – – 36.09 1.3 161
Ž4–8. 16.38 Ž1.2–3.5. 4.5 Ž1.19–2.69. 6.62 Ž5.33–4.33. 5.99 Ž4.64–3.15. 9.43 Ž12–10.50. 7.64 Ž13.83–11.83. 0.39 Ž9.25–11.83. 11.00 Ž0.2–1.0. 4.55 0.35 Ž0.54–0.74. 1.20 0.68 Ž0.17–0.27. 0.66 Ž0.1–1.5. 9.75 Ž0–0.5. 6.53 Ž0.86–2.5. 7.17 Ž2.5–2.5. 4.84 Ž4.35–4.35. 2.06 Ž2.91–2.91. 0.46 Ž0.73–0.73. 2.57 Ž0.13–1.9. 1.80 Ž11.82–11.32. 0 Ž3.34–4.32. 0 Ž15.52–8.5. 1.95 Ž12.5–8.5. 0.20 Ž7.10–6.60. 8.16 Ž0.2–1.0. 3.50 Ž7.5–5.50. 3.93 Ž0.7–2.5. 1.20 2.71 1.82 19.86 2.4 150
Ž13.38–18.00. 15.77 Ž3.00–5.00. 4.5 Ž5.00–6.62. 6.17 Ž8.99–5.49. 7.49 Ž9.43–8.43. 10.97 Ž7.64–7.64. 3.89 Ž0.39–0.39. 1.69 Ž11.00–11.50. 11.82 Ž4.0–5.09. 4.59 Ž0.25–0.35. 0.38 Ž1.20–1.20. 1.10 Ž0.50–0.68. 0.78 Ž0.56–0.66. 0.20 Ž9.20–11.00. 11.27 Ž5.53–7.00. 8.87 Ž6.17–7.60. 5.68 Ž4.84–6.53. 3.83 Ž2.06–2.06. 2.26 Ž0.46–0.46. 0.46 Ž2.57–2.57. 3.01 Ž1.00–2.00. 2.67 0 0 Ž1.95–1.75. 2.22 Ž0.50–0.10. 0.20 Ž6.01–5.61. 8.70 Ž3.00–4.50. 4.71 Ž4.93–2.93. 3.03 Ž1.20–1.20. 1.57 Ž1.50–3.00. 1.08 Ž1.82–1.82. 2.30 15.91 2.7 149
Ž13.0–16.5. 12.47 Ž3.00–5.00. 2.96 Ž4.50–6.17. 6.17 Ž9.49–6.99. 10.85 Ž10.97–9.97. 6.98 Ž4.89–3.89. 2.79 Ž0.39–2.00. 11.73 Ž11.82–10.82. 8.45 Ž4.0–5.09. 3.45 Ž0.35–0.38. 0.38 Ž1.10–1.10. 0.96 Ž0.50–0.78. 0.41 Ž0.40–0.20. 0.2 Ž11.00–12.5. 10.82 Ž7.87–9.00. 5.7 Ž6.68–5.68. 6.26 Ž3.83–4.83. 3.22 Ž1.26–2.26. 2.03 Ž0.46–1.46. 1.18 Ž3.01–3.33. 3.33 Ž2.67–2.67. 4.49 0 0 Ž2.22–2.00. 2.08 Ž0.50–0.10. 2.42 Ž7.20–6.80. 7.72 Ž4.81–5.71. 4.87 Ž6.30–4.03. 3.39 Ž1.57–1.57. 1.96 Ž1.08–1.08. – Ž2.35–2.35. 2.43 15.91 2.65 149.85
J.M. Adams, W.M. Post r Global and Planetary Change 20 (1999) 243–256
tropical forest biome, due to the relative abundance of a range of forms of palaeoevidence from the temperate regions. For our ‘low forest’ scenario we suggest even greater reductions in forest area than is considered most likely by Adams and Faure, and more in line with the views of authors such as Frenzel Ž1968. and Clapperton Ž1993..
3. Results and discussion If for consideration one simply compares the range of possible values ŽFig. 2. which would result from a "30% error limit on the average soil carbonate content in each biome, the glacial-to-Holocene change could in principle involve no change or even a slight increase in soil carbonate, or as much as a 1000 Gt decrease. However, if due to sampling error a particular biome in the present-day world has Žfor example. a 30% lower soil carbonate content on average than the preferred figure which Batjes and Sombroek Ž1997. give, this error will presumably follow in the same direction for both glacial and interglacial conditions, rather than making it possible to compare an extreme q30% error for one timeslice against a y30% error for another timeslice. This constraint will tend to ‘pull in’ the true limits of uncertainty on the glacial-to-Holocene change in soil carbonate carbon. At a more general level, the width of error limits should not be allowed to distract from the fact that a several hundred Gt decrease in carbonate carbon remains much more likely from the general biome distributions, and from what is known of their soil carbonate content from present-day nearest analogs. A comparison of the LGM timeslice with the three Holocene timeslices ŽTable 3, Fig. 2. suggests that a major Žaround 40–50%. decrease in soil carbonate carbon occurred between LGM and early-tomid Holocene conditions, due to the greatly decreased areas of arid and semi-arid biomesrclimate zones ŽTable 3.. For example, if one takes the ‘preferred’ figure for the LGM, and compares it to the preferred figures for the Holocene ŽFig. 2., the shift is towards about 400–500 Gt less carbon. As an uppermost estimate for the shift, subtracting the ‘low forestrdry world’ q30% error scenario for the LGM Ž1497 GtC. from a ‘high forest, wet world’ q30%
251
error scenario for the 8000 14C ya, scenario Ž731 GtC. gives a larger shift of 766 GtC. As a lowermost estimate, a shift from the y30% error ‘high forestrwet world’ scenario for the LGM Ž697 GtC. to the ‘low forestrdry world’ y30% error scenario for the early Holocene Ž468 GtC., would give a relatively low shift of 229 GtC. Thus, the overall possible range for the ‘true’ figure of the inorganic carbon storage change can be seen as lying somewhere between 766 and 229 Gt, with the figure most likely to have been around 400–500 Gt considering the per-unit-area estimates given by Batjes and Sombroek, and the approximate consensus global biome distributions produced by QEN ŽFig. 1a–d.. This shift to less carbon in soil carbonates with the onset of interglacial conditions would have been associated with a major efflux of soil carbonate, most of it presumably chemically weathered in the following reaction, partly catalysed by biological activity in the soil; calcium carbonateq carbon dioxideq water | dissolved calcium ions q dissolved bicarbonate ions CaCO 3 q CO 2 q H 2 O | Ca2qq 2 Ž HCO 3 .
y
Smaller amounts of magnesium carbonate present in soil carbonates would have been dissolved in the same manner, and a minor fraction Žjudging from the ratio of sulphate to carbonate in present-day groundwater and river water compositions; Probst, 1994. weathered by sulphuric acids to give dissolved sulphate salts. This latter reaction would release a small proportion of CO 2 to the atmosphere. In the present-day world, weathered bicarbonate enters groundwater or is washed down in surface flow to rivers, and ultimately the oceans ŽProbst, 1994.. In order to be weathered into dissolved bicarbonate, the diminishing reservoir of soil carbonate would have to take up an additional 400–500 Žouter limits about 750–200. GtC as CO 2 derived from oceanratmosphere sources. If 2.1 GtC equals approximately 1 ppm CO 2 , this represents the equivalent of 190–238 ppm of CO 2 in the atmosphere. The soil organic carbon reservoir under LGM and late glacial conditions also seems to have been much less
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Fig. 2. Calculated global GtC in soil carbonates. Upper and lower limits are shown for ‘low forest’ and ‘high forest’ vegetation distributions, respectively. Outermost limits allow for uncertainty in mean soil carbonate content in each life zone. Previously published present-day estimates by Schlesinger and by Batjes are also shown for comparison.
than during the Holocene and would thus have been unable to supply this extra quantity of carbon by oxidation ŽAdams and Faure, 1998.. As Batjes and Sombroek Ž1997. point out, the detailed dynamics of soil carbonate accumulation and loss during a future switchover in climate and water regime are highly uncertain. The same can be said of past changes, and our calculations here as-
sume that a rough equilibrium state had been achieved given the climate and vegetation conditions at each timeslice—an assumption which may be difficult to test given the present state of knowledge of soil carbonate dynamics, but which might not be too inaccurate considering the thousands of years for which the general climatic and vegetation conditions at each timeslice persisted.
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Table 3 Calculated global soil carbon storage in carbonate. Values are in t = 10 8. Preferred scenario is outside the parentheses, while the first value in parentheses is for the ‘wetter climate’ scenario with more forest, the second is for the ‘drier climate’ scenario with less forest and more arid and semi-arid vegetation Ecosystem code
LGM
8000 years
5000 years
Present pot.
Ž1. TrRfr Ž2. MsnFor Ž3. TrWd Ž4. TrThnSc Ž5. TrSDes Ž6. TrGrs Ž7. TrDes Ž9. Sav Ž10. WarmTemFor Ž11. CtGiantConfr Ž12. TrMontFor Ž13a. MdtnForest Ž13b. MdtnScrub Ž14. CoolTempFor Ž15. SouthTaiga Ž16. Mid Tg Ž17. OpnBWd Ž18a. Temp wdld Ž18b. Temp Scrub Ž19a. Mtnrdry Tu Ž19b. Lowl Tu Ž20a. Steppic St-Tu Ž20b. Tundric St-Tu Ž21. PolarrMontane Desert Ž22. Temperate Desert Ž23. Temperate Semi-Desert Ž24a. Moist Steppe Ž24b. Dry Steppe Ž25. Forest Steppe Ž26. Forest Tundra Ž27. BogrSwamp Ž28. Ice Ž29. Lakerriver Total carbon
27.8 Ž37.5–20. 13.5 Ž24.5–8.4. 35.7 Ž80.7–35.7. 266.4 Ž346.4–426.4. 371.2 Ž252–371.2. 840 Ž840–960. 1659.6 Ž1419–1740. 980 Ž946–740. 11.4 Ž11.4–740. 0 8.1 Ž11.1–7.5. 0 5.1 Ž8.1–5.1. 6 Ž45–3. 0 24 Ž50–17.2. 37.5 Ž52.5–22.5. 130.5 Ž130.5–130.5. 291 Ž291–291. 7.3 Ž7.3–7.3. 1.3 Ž19–1.3. 1182 Ž1182–1182. 334 Ž432–334. 1940 Ž1190–1940. 1562.5 Ž1062–1562. 710 Ž810–910. 50 Ž100–20. 650 Ž550–750. 71.5 Ž71.5–38.5. 0 0 0 0 11,216
81.9 Ž90–66.9. 31.5 Ž35–21. 198.1 Ž198.5–150. 479 Ž439–639.2. 754 Ž674–834.4. 611.2 Ž611–611.2. 46.8 Ž14.4–46.4. 869 Ž920–880. 68.2 Ž68.2–60. 5.25 Ž5.2–3.75. 18 Ž22.5–18. 20.4 Ž23.4–15. 19.8 Ž16.8–16.8. 292.5 Ž330–276. 130.6 Ž140–110.6. 143.4 Ž152–123.4. 72.6 Ž78.4–87.6. 61.8 Ž61.8–91.8. 46 Ž46–446. 25.7 Ž15.7–25.7. 18 Ž25–10. 0 0 243.7 Ž218.7–243.7. 25 Ž12.5–62.5. 816 Ž461–601. 350 Ž400–300. 393 Ž293–693. 66 Ž121–121. 149.0 Ž165–137.5. 0 0 0 6,038
78.8 Ž81–65. 31.5 Ž38.5–22.4. 185.6 Ž210.3–135.2. 479.2 Ž559.5–839.6. 754.4 Ž717.6–877.6. 611.2 Ž311.2–471.2. 46.8 Ž46.8–46.8. 865.6 Ž705.6–945.6. 68.25 Ž76.3–60. 5.7 Ž5.7–5.25. 18 Ž37.5–30. 23.4 Ž23.4–15. 6 Ž6–12. 338.1 Ž375–330. 177.6 Ž180–157.4. 113.6 Ž100–133.6. 60.15 Ž60–57.4. 67.8 Ž67.8–37.8. 46 Ž50–246. 33.1 Ž33.3–33. 26.7 Ž26.7–21. 0 0 277.6 Ž250–277.6. 25 Ž12.5–62.5. 870 Ž600–820. 471 Ž671–471. 303 Ž603–550. 86.3 Ž86.3–86.35. 59.4 Ž86.9–49.5. 0 0 0 6,086
62.3 20.7 185.1 868.2 558.6 223.2 1407 676.6 51.8 5.7 14.5 12.4 6 324.1 114.4 125.6 48.1 60.8 118 33.3 44.9 0 0 260 302.5 772 487 339 107.8 0 0 0 0 7,229
3.1. Other interacting carbonate sinks r sources It is unclear how much of the ‘missing’ 200–750 Gt carbon taken up by postglacial soil carbonate weathering is still held in the oceans or in groundwater, and what proportion of it cycled rapidly through the system, depositing carbonate on sea floors or coral reefs and re-entering the atmosphere. The diminution of the soil carbonate reservoir adds to the problems involved with modelling the glacial carbon cycle, which already cannot account for the source of CO 2 needed to replenish organic soil and vegetation
stocks on land during the interglacial ŽAdams and Faure, 1998.. One possibility is that the diminishing caliche reservoir on land has been approximately balanced by the accumulating coral reef reservoir during the Holocene: an estimate by Faure et al. Ž1988. suggests that about 500 Gt inorganic C has accumulated in coral reefs since the start of the Holocene Žallowing for 100 Gt internally recycled carbonate due to bioerosion; H. Faure, pers. commun., 1998.. Since coral reefs grow by splitting dissolved bicarbonate to release insoluble calcium carbonate and CO 2 Žused
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in photosynthesis by coralline algae., this mechanism could approximately balance the ‘extra’ carbon taken up in caliche weathering on land. The CO 2 released may partly have been retained as organic matter within the buried coral, but after the death of the coral organism most of it would presumably have been released as CO 2 into the general ocean–atmosphere system Žgiven the fairly low organic matter content of coral rock.. Hence, overall, there may be little or no net flux of atmospheric CO 2 into the ocean’s dissolved carbonate reservoir, with calcretederived bicarbonate merely acting as a temporary ‘ vehicle’ for CO 2 . In addition to corals, there is the quantity of carbonates precipitating in muds and oozes on the continental shelves Žwhich increased in area as the sea level rose during deglaciation. and the continental slopes, although no specific estimate of this accumulation since the beginning of the Holocene is available. However, during the early Holocene, the rapid initial loss of calcrete might not have been fully counterbalanced by the growth rate of reefs, and the slowness of sea level rise may have delayed the main episodes of abiotic shelf carbonate precipitation until well into the Holocene. Other subaerial bicarbonate sources need to be considered in addition to soil carbonates, to place our estimates in their true context. The other sources seem likely to have changed in the same direction as the calcretes, taking up more CO 2 and sending more bicarbonate to the oceans under interglacial conditions. It is probable that the quantity of carbon taken up each year by silicate weathering Žalso ultimately the source of soil carbonate. was greatly affected by the glacial-to-interglacial climate and vegetation change ŽAdams, 1993, 1995; Adams and Faure, 1997. perhaps adding several further hundreds of gigatonnes to the cumulative extra amount of CO 2 taken up by the land surface, as opposed to any corresponding time interval of the last glacial Oxygen Isotope Stage 2. This would further have added to the burden of dissolved carbonate entering the oceans, whose fate needs to be explained. Karst areas of limestone on land weather rapidly and provide a further major sink of CO 2 and source of dissolved carbonate to the oceans, which would have varied between glacial and interglacial conditions with the large changes in moisture conditions.
Present-day uptake of carbon by karst carbonate weathering may be of the order of 0.2–0.6 GtCryr ŽIGBP, 1998.. As Ludwig et al. Ž1998. have suggested, carbonate weathering influx Žand rates of CO 2 uptake as bicarbonate leaving for the ocean. on dead reefs on exposed continental shelves during the Last Glacial might have partly offset the change resulting from the switch from dry to moist climates around the world. However, it is necessary to bear in mind that sea levels took much longer to rise than climate took to change; large areas of coral reef would still have been exposed to subaerial weathering under the renewed warmer and moister climates of the interglacial. This itself should have provided a great increase in flux of carbonate to the oceans between glacial and interglacial conditions; an additional flux which has not been considered in this paper and which awaits further work. 3.2. The context of the atmospheric CO2 and ocean carbonate record Interpretation of the ocean carbonate Žlysocline. record ŽBroecker and Sanyal, 1997., needs to take into account the range of different processes including calcrete weathering, which may have produced a rapid increase in efflux Žof the order of hundreds of gigatonnes. of carbon in dissolved bicarbonate to the oceans during the several thousand year period of deglaciation. The role of groundwater accumulation of bicarbonate and re-deposition of carbonate deeper within the soil system needs to be considered, but it is unlikely that this can account for more than a small fraction of the loss of carbonate from calcretes. The general record of the balance between carbonate preservation and dissolution in ocean sediments suggests that during deglaciation there was an increase in carbonate preservation at greater depths Ža deepening of the lysocline, the level below which carbonate redissolves.. This is consistent with withdrawal of CO 2 from the ocean by a regenerating land biota ŽBroecker and Sanyal, 1997., but it may also partly be a product of more carbonate from weathering being ‘shuttled’ to the oceans as bicarbonate and released by carbonate precipitation, with the CO 2 from this dissociation ultimately being absorbed as ecosystem organic carbon storage back on
J.M. Adams, W.M. Post r Global and Planetary Change 20 (1999) 243–256
land. It is important to bear in mind, however, that variations in ocean primary productivity and in watermass circulation patterns could also be important in controlling changes in the lysocline. Furthermore, the resulting flux—involving hundreds of Gt extra CO 2 transport by bicarbonate during a several thousand year period—needs to be incorporated into carbon cycle models. A decrease of 750–200 GtC in soil calcrete would require some 750–200 GtC of CO 2 to dissolve it into bicarbonate. This soil calcrete flux alone is equivalent to between 357 and 95 ppm atmospheric CO 2 Žif one assumes a relationship of 2.1 GtC s 1 ppm CO 2 .. This is enough to have had a significant effect on the postglacial accumulation of CO 2 in the atmosphere if there was a delay in the recycling of this CO 2 back out of the oceans taking more than a few hundred years. To this must be added the effect of a probable major increase in subaerial weathering of exposed reefs, older carbonate rocks and silicate rocks as climates around the world switched to their interglacial state.
Acknowledgements The advice of Hugues Faure, Ellen Thomas, Lloyd Keigwin and our two anonymous reviewers is much appreciated. Jonathan Adams was supported by the Lockheed-Martin on a Eugene P. Wigner Fellowship Program at Oak Ridge National Laboratory. Research sponsored by the U.S. Department of Energy, Carbon Dioxide Research Program, Environmental Sciences Division, Office of Biological and Environmental Research, under contract DE-AC05-96OR22464 with Lockheed Martin Energy Research Corp.
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