Biogeochemical cycles I carbon and oxygen

1. https://webfiles.uci.edu/setrumbo/public/IMPRS/Trumbore_IMPRS2.pdfhttps://webfiles.uci.edu/setrumbo/public/IMPRS/Trumbore_IMPRS2.pdf https://webfiles.uci.edu/setrumbo/public/IMPRS/Trumbore_IMPRS3.pdf Biogeochemical cycles I carbon and oxygen 9:00 – 10:30 Wednesday Dec 8

2. Carbon Cycle Questions What caused changes in CO2 between glacial and interglacial? Did CO2 force or respond to climate? More on isotopes and what they can tell you More detail on the ocean biogeochemical cycling What is the long-term fate of CO2 we add to the atmosphere?

3. Controls on CO2 vary with timescale • Millions of years; volcanic CO2 supply and weathering uptake (tectonics) • Thousands of years (glacial/interglacial cycles) Ocean determines atmospheric CO2 • Decades-Centuries: many processes (growth/decomposition in forest, soil OM exchange, changes in ocean circulation, ….. • Seasonal cycle • Ocean gas exchange • Terrestrial biosphere

4. Distribution of Carbon; 1015 grams = 1 Petagram (Pg) ATMOSPHERIC CO2 640 X1015 g C Response times are seasons to centuries LIVING BIOMASS 830 X1015 g C DISSOLVED ORGANICS 1500 X1015 g C Response times are centuries to millennia ORGANIC CARBON IN SEDIMENTS AND SOILS 3500 X1015 g C CO2 DISSOLVED IN OCEANS 38,000 X1015 g C LIMESTONE AND SEDIMENT CARBONATES 18,000,000 X1015g C Response times are tens of thousands to millions of years TRAPPED ORGANIC CARBON: NATURAL GAS, COAL PETROLEUM, BITUMEN, KEROGEN 25,000,000 X1015 g C

5. Long term control- balance of weathering rate and CO2

6. Last Glacial Maximum (LGM) • Time of maximum ice sheet extent centered on 21 ka (19-23 ka) • Glacial world was significantly different from today: • Ice sheets/sea level • Temperature • Greenhouse gases • Aridity • Winds • Vegetation • Ocean circulation • Continental configuration & insolation were nearly identical to today… • pCO2 and ice volume are most likely factors affecting LGM climate • Major boundary conditions are well known and abundant paleoclimate data is available – a crucial test for climate models!

7. Where did the C go during glacials? • pCO2 changes from 190-280 ppm (30%) in a few 1000 y, this cannot be due to weathering or volcanic CO2. • Fast changes during Quaternary can only be explained by rapid C exchange among surface reservoirs! A useful way to measure how C has moved among various reservoirs is using δ13C Total change of atmospheric C: 90 ppm or nearly 200 PgC

8. How to get C into the deep ocean? • Physical changes (Solubility pump) • Temperature & salinity • Isolation of deep from surface waters (decreased ventilation) • Stronger biological pump • Fe fertilization • Increase of whole ocean nutrient content • Change in Redfield ratios (more efficient C pumping) • Changes in ocean [CO32-] • Increased CaCO3 weathering • Decreased coral reef growth • Change in C0rg:CaCO3 export to deep ocean

9. The solubility and biological pumpsChanges in ocean circulation Surface waters equilibrate quickly; CO2 reacts with water Falling particles move organic carbon into the deep ocean Sinking waters in polar regions isolate water that has equilibrated at the surface (cold waters)

10. Temperature/Salinity mechanism is “easy” to test… • Cooler SSTs = increased CO2solubility; lowers CO2 by 30 ppm • But, also higher salinity, decreases Solubility (+6.5) Net T & S effect  -23.5 ppm (as opposed to 90 ppm observed) … not enough

11. Was it stored on land? Crowley et al., 1997, JGR

12. Was it stored on land?Still unresolved, but in evidence so far is that LESS C was stored on land…Where did it go? The deep ocean is the only reservoir big enough and slow-exchaging enough • Decreased temperate forests • Increased northern tundra • Decreased tropical rain forests • Reduced growth due to low pCO2 Crowley et al., 1997, JGR

13. 13C changes in benthic foraminifera should show this transfer The 13C in benthic forams varied between the last glacial and today http://www2.ocean.washington.edu/oc540/lec01-28/ Comparison of the d13C records from equatorial (V19-30) and northeast Pacific (W8709A-8) cores spanning the last glacial cycle. Based on this record, the glacial ocean 13C was roughly 0.4per mil  lighter during the LGM (indicating transfer of isotopically light C from land to ocean), and consistent with a smaller land biosphere. However, the decrease predicted by transferring 530 PgC is less, only -0.35 per mil; something else going on…

14. Remember:  Land pants (C3) have  13C of about -25 per mil  (R =0.975= 13C/1000 +1))          Ocean total CO2 (Holocene)  13C is about +0.50 per mil  (R=1.005)       LGM Ocean total CO2    = 0.50  (Holocene value) minus 0.35 per mil = 0.15 per mil We can use the difference in 13C between ocean+atmosphere today and in the LGM to estimate the how much less land C there was on the LGM by mass balance: Carbon mass balance: Land]today  + OA]today = Land]glacial + OA]glacial 2,000 (land today)   + 3,6500 (35,100 in ocean, 500 in preindustrial atmosphere)  = Total = 38,600 Pg of carbon 13C mass balance: (2000)(0.975)  +  (38,600)(1.0005) =  Land]glacial (0.975)  + OA]glacial(1.00015)                                            = Land]glacial (0.975)  + [38,600 - Land]glacial](1.00015) Solving for Land]glacial we get ~1500 Pg C  (or 500 Pg C less than today) Other differences:                    Preindustrial                 LGM Land :                2000                         1500     (from 13C in benthic forams)Atmosphere         500                          360     (from pCO2 in ice cores)Ocean             31,500                      35,740  (by difference)

15. NOTE: There are some problems here.  • The 500 Pg C difference between LGM and today in the biosphere calculated using 13C change is at the very low end of the range that has been estimated from paleovegetation maps (700-1300 PgC) • There are a number of potential problems with 13C in forams, mostly involved with • differences in 13C between coexisting benthic species (vital effects) coupled with selective dissolution • the tendency of benthic forams to use DIC that is in part derived from the decomposition of organic material in sediment pore waters. • (3) the distribution of C3 and C4 plants in the LGM was likely different (i.e. if C3 biomes were replaced with C4 vegetation, there in theory be a shift in 13C isotopes without a shift in biomass on land). • (4) the 13C record differs from one area of the ocean to the next - this likely reflects changes in paleo-ocean circulation/ biological pump (more on this later).

16. Carbon species in seawater Dissolved CO2 pCO2 (or as it is more correctly expressed [H2CO3] ) is a minor constituent of seawater carbon ~1% Bicarbonate ion (HCO3-) is ~90% of the carbon at ocean pH (8.2)Carbonate ion  (CO32-) is ~10% of the total carbon Total Dissolved Inorganic C (TDIC) =H2CO3 + HCO3- + CO32- Alkalinity (ALK) is the excess of cations over weak acid anionsIn seawater, and ignoring borate for the moment, ALK is proportional to HCO3- + 2CO32- Therefore, carbonate ion may sometimes be approximated as ALK - TDIC (in surface water) The major chemical equilibrium we deal with is:CO2 +CO32- +H2O <==>2HCO3- The equilibrium constant, varies with temperature and salinity (and pressure)

17. TDIC (= *H2CO3   + HCO3- + CO32- )is influenced by three processes: (1) CO2 exchange with the atmosphere (2) photosynthesis/respiration (3) carbonate precipitation and dissolution Alkalinity (Charge balance ~ HCO3- + 2CO32-) is influenced by: (1) carbonate precipitation and dissolution (2) organic matter formation and decomposition (a small amount, through NO3- uptake and release)

18. Seawater DIC is primarily HCO3- and CO32- CO2(aq) increases at lower pH

19. Revelle Factor CO2 increases by ~10% when DIC increases by ~1% Low latitudes have Higher CO32- And lower R factor What does this mean? CO2 +CO32- <==>2HCO3- Increasing CO2 drives the reaction to the right, reducing CO32- but making more HCO3- There is a lot of DIC in the ocean, converting one form to another does not change the total amount much; relative change is small

20. WHAT WILL BE THE IMPACT ON OCEAN CHEMISTRY AND ATMOSPHERIC CO2? The change in land carbon actually added carbon to the atmosphere in the LGM;  some of that CO2 would dissolve immediately in the surface ocean, and ultimately be reflected in increased CO2 in deep waters.  The increased CO2 would cause dissolution of carbonates in the deep sea (over a timescales of thousands of years). DEEP WATER CHANGES IN CARBONATE CHEMISTRY                           Interglacial Ocean         LGM                             LGM                                                          (before Calcite)      (after calcite dissolution)Alkalinity                2270 (meq/kg)           2270                    2322 (2270 + 52)Total CO2  (TDIC)   2085 (mmol/kg)         2115                     2141 (2115 + 26)CO32-                      129 (mmol/kg)            112129pCO2                       280 (matm)                 336                             296 DpCO2                                               +56                             +16 Adding or removing CO2 does not change alkalinity much (why not?) 500/35,600 is a 0.14% increase in atmosphere/ocean C –How much goes into the ocean (vs. atmosphere) depends on the Revellefactor. Adding a 500 Pg CO2 means about a 50 ppm rise in CO2 (with RF of 0.1) Because the CO32-   is lower, the deep waters are undersaturated and CaCCO32-    will dissolve until equilibrium is re-established.

21. If we add 500 Pg C to the atmopshere, how much will by the surface ocean and how much will remain in the atmosphere? Revelle factor (DpCO2/pCO2)/(DDIC/DIC) ~10 If you equilibrate with just the surface ocean (~1020 PgC) DpCO2 = pCO2* 10 *(DDIC/DIC); DpCO2 = 6(DDIC) For the deep ocean (38,000 PgC = DIC); DpCO2 = 0.11DDIC But mass balance says DDIC = 500PgC – DpCO2 So for pCO2 = 480 (LGM) and DIC = 1020; DpCO2 (1+1/6) = 500; DpCO2 = 430 PgC For DIC = 38,000 (i.e. equilibrate with whole ocean), DpCO2 (1+1/.11) = 500; DpCO2 = 50 PgC

22. Negative feedback – precipitation rate of CaCO2 in the ocean (the depth of the lysocline). Buffers changes in deep ocean CO3-- Solubility Ksp = [Ca+2][CO32-]; Ksp is dependent on pressure, temperature (increases with pressure – so that carbonate formed in the surface ocean will dissolve at depth) Le Chatlier’s rule – if you decrease[CO32-] in deep water in contact (equilibrium) with CaCO3 in sediments, you will dissolve carbonate until equilibrium is reestablished)

23. The bottom line:  A smaller biosphere in the LGM means HIGHER CO2 (by about 16 ppm if the biosphere lost 500 PgC to the atmosphere/ocean).  An even smaller biosphere (as has been proposed by those making estimates from paleoecology) means an even higher LGM pCO2) SUMMARY WITH TEMPERATURE/SALINITY CHANGES: Terrestrial C decrease                   +15 ppmOcean cooling                               -30  ppmOcean salinity increase                 +6.5 ppm Total                                              -8.5 ppm SOMETHING ELSE IS NEEDED TO EXPLAIN GLACIAL-INTERGLACIAL CO2 CHANGE!

24. Biological ‘pump’ • 12C preferentially taken up by phytoplankton • surface waters (and shells) enriched in 13C 12C enriched from oxidation of organic matter

25. d13C of DIC in seawater Surface water Photosynthesis preferentially removes 12C, leaves behind water enriched in 13C Deep water – also along ‘conveyor’ Remineralization of organic matter adds 12C enriched material, lowering d13C Ocean 13C Efiiciency of the biological pump can be reflected in the difference in 13C between surface and deep water. There is therefore (or should be) a relationship between 13C and CO32- ion content of deep water

26. Possible mechanism: Increased nutrient utilization (or supply)?

27. A proxy for the biological pump? • Surface – deep water d13C (preserved in foram shells) is a measure of the strength of the biological pump • Glacial periods = Larger difference = stronger pump • More C stored in deep sea • But some problems: • Other sources of d13C variability • Foram d13C is complicated… • Increased C pumping should decrease deep ocean [CO3]2-, but no evidence for shallower lysocline

28. Ocean circulation at the LGM • Changes in Atlantic circulation have been linked to past climate changes (glacial-interglacial and abrupt) • In modern Atlantic , a net oceanic heat transport from North to South. If we perturb this transport, we alter climate Modern ocean circulation can be visualized using Wally Broecker’s ocean conveyor…

29. Bomb 14C SS NCP SOce 14C in DIC and DOC in the Deep Conveyor A measure of the time since deep water equilibrated with the atmosphere -525 to -390‰ Williams and Druffel, 1987; Bauer et al. 1992; Druffel and Bauer, 2000

30. The ‘age’ of carbon increases from deep Atlantic to deep Pacific (this is where the ‘conveyor’ idea came from) 2050 – 670 = 1380 yr 5980 – 3160 = 2820 yr

31. Possible mechanisms… • Stronger overturning of Antarctic intermediate waters could have delivered more nutrients to surface waters & increased biological pump • Polar alkalinity hypothesis **Remember: CO2 + CO32- + H2O  2HCO3- • Today: NADW dissolves little CaCO3 and upwells in S. Ocean with low [CO32- ],leaving S. ocean surface waters (and overlying atmosphere) with high CO2 • Glacial: Southern source waters with high CO2 (more corrosive) expanded , dissolved more CaCO3 ,and returned more CO32- to Antarctic surface waters. • Broecker and Peng, 1989 proposed that this could explain ~ 40 ppm decrease in atmospheric CO2, , but more recent sediment data does not support this…

32. It is likely that the carbonate system plays an important role though… • pCO2 in surface water is a function of both DIC & Alk • Changes in mean inventory of either would impact surface water, and hence, atmospheric pCO2 = HCO3- + 2CO32- + OH- - H+ … = CO2(aq) + H2CO3 + HCO3- + CO32-

33. The answer likely lies in the Southern Ocean • Co-evolution of Antarctic temperature & atmospheric CO2 • Nutrients are currently underutilized • Southern ocean ventilates large volumes of ocean interior • Two mechanisms for changes in S. ocean nutrient utilization: • Physical changes could isolate deep waters from surface, limiting CO2 degassing • Biological changes due to increased Fe (and Si?) fertilization by dust (increased Corg:CaCO3 export)

34. Summary • It is likely that glacial-interglacial CO2 changes require a variety of mechanisms to explain. • The current frontrunners include: • T & S changes (-20 to 30 ppm) • Southern ocean mechanisms (major contributor) • Certain mechanisms (i.e. changes in whole ocean [CO32-] )seem unlikely due to disagreement with available proxy data (which is admittedly scarce) • Much work remains to be done to resolve this!

35. Carbon Cycle Part II What is the fate of CO2 we add to the atmosphere by fossil fuel burning and land use?

36. http://scrippsco2.ucsd.edu/graphics_gallery

37. http://www.esrl.noaa.gov/gmd/obop/mlo/programs/esrl/ccg/img/img_global_co2.jpghttp://www.esrl.noaa.gov/gmd/obop/mlo/programs/esrl/ccg/img/img_global_co2.jpg

38. Where does the other ~40% go??? Also, what happens to CO2 from deforestation (not counted here) Source: Ralph Keeling, SIO

39. Deforestation: Clearing of forests (formerly in the US, now in the tropics) Responsible for ~40% of total C emissions since 1850 In 1990s 0.5 to 2 GtC/year (8-25% of total emissions)

40. 2000-2009 (PgC) 10 5 Source deforestation CO2 flux(PgC y-1) 1.1±0.7 5 Sink 10 1950 2000 1900 1850 Time (y) Human Perturbation of the Global Carbon Budget Global Carbon Project 2010; Updated from Le Quéré et al. 2009, Nature Geoscience; Canadell et al. 2007, PNAS

41. 1000 Latin America 800 S & SE Asia Tropical Africa 600 400 CO2 emissions (Tg C y-1) 200 0 -200 1850 1860 1870 1880 1890 1900 1910 1920 1930 1940 1950 1960 1970 1980 1990 2000 2010 Time (y) Emissions from Land Use Change (2000-2009) R.A. Houghton 2010, personal communication; GFRA 2010

42. Global Fire Emissions Database (GFED) version 3.1 1400 1200 1000 Fire Emissions from deforestation zones(Tg C y-1) America 800 Africa Asia 600 Pan-tropics 400 200 0 01 99 1997 07 05 2003 2009 Time (y) Fire Emissions from Deforestation Zones van der Werf et al. 2010, Atmospheric Chemistry and Physics Discussions

43. Use of remote sensing to determine area deforested leads to reduced estimates of CO2 emissions Estimates for the 1990’s Ref. 106 ha a-1PgC a-1 Houghton (FAO) 15.5 2.2(±0.8) DeFries 5.6 0.9(±0.4) E Van der Werf et al. 2009 Nature Geoscience

44. 2000-2009 (PgC) 10 fossil fuel emissions 7.7±0.5 5 Source deforestation CO2 flux(PgC y-1) 1.1±0.7 5 Sink 10 1950 2000 1900 1850 Time (y) Human Perturbation of the Global Carbon Budget Global Carbon Project 2010; Updated from Le Quéré et al. 2009, Nature Geoscience; Canadell et al. 2007, PNAS

45. 2000-2009 (PgC) 10 fossil fuel emissions 7.7±0.5 5 Source deforestation CO2 flux(PgC y-1) 1.1±0.7 5 Sink 10 1950 2000 1900 1850 Human Perturbation of the Global Carbon Budget Time (y) Global Carbon Project 2010; Updated from Le Quéré et al. 2009, Nature Geoscience; Canadell et al. 2007, PNAS

46. 2000-2009 (PgC) 10 fossil fuel emissions 7.7±0.5 5 Source deforestation CO2 flux(PgC y-1) 1.1±0.7 atmospheric CO2 4.1±0.1 5 Sink 10 1950 2000 1900 1850 Time (y) Human Perturbation of the Global Carbon Budget Global Carbon Project 2010; Updated from Le Quéré et al. 2009, Nature Geoscience; Canadell et al. 2007, PNAS

48. Suess Effect: Fossil fuel-driven depletion of atmospheric D14C SUESS HERADIOCARBON CONCENTRATION IN MODERN WOOD, SCIENCE, 122 (3166): 415-417 1955 Jacobson [2000]

49. Fossil fuel has d13C of -21 to -27 per mil If all emissions are taken up by the biosphere, the d13C of atmospheric CO2 should not change. Dissolution in the ocean preferentially removes 12C more than 13C, so we would expect a decline in 13C of atmospheric CO2 http://scrippsco2.ucsd.edu/graphics_gallery

50. Land/Ocean sinks from 13C • The basic equation C3 ~ 20‰C4 ~ 4.4‰O ~ 2‰ • A terrestrial sink makes the atmosphere heavier ( more enriched in d13C) • An ocean sink has little effect on atmospheric 13C • A C4 sink looks like ocean to the atmosphere • As with CO2, the atmosphere shows signs of terrestrial uptake, mainly at high latitudes. • But the “disequilibrium” problem makes the interpretation of 13C very challenging.