https://webfiles.uci/setrumbo/public/IMPRS/Trumbore_IMPRS_1

1. https://webfiles.uci.edu/setrumbo/public/IMPRS/Trumbore_IMPRS_1.ppthttps://webfiles.uci.edu/setrumbo/public/IMPRS/Trumbore_IMPRS_1.ppt

2. Earth is an Open System for Energy Some Energy is reflected to space Some Energy is absorbed and heats the earth; this drives biogeochemical cycles Earth intercepts energy radiated by the sun A comparatively very small flux of heat from radioactive decay of elements within earth – drives lithosphere cycles Earth has temperature, so emits radiation (energy) to space

3. Earth is a closed system for all elements heavier than helium Implications: Finite availability of resources Changes propagate through system Biogeochemical cycles – movement of elements between litho-, hydro-, bio-, atmosphere

4. Tools and some basics • BiogeochemistryScientific study of the chemical, physical, geological, and biological processes and reactions that govern the composition of the natural environment, and the cycles of matter and energy that transport the Earth's chemical components in time and space. (from wikipedia) • Global Biogeochemical CyclesDescriptions of the movement of elements (e.g. C, N, S, O, P) among global spheres (biosphere, hydrosphere, pedosphere, atmosphere, lithosphere, antrhoposphere) • Why study them?Humans have altered most cycles in ways that are having global importance (at least on timescales of the next few millennia); biogeochemical cycles control the composition and longterm evolution of Earth’s atmosphere and climate

5. Abundance of elements in solar system depends on nucleosynthesis H, He are most abundant (not shown here) (from big bang, fusion of H to form He in stars) Elements like 12C, 16O, 20Ne up to 56Fe formed by burning He in massive stars (burning releases energy) Elements heavier than 56Fe are formed in supernovae A few elements (cosmogenic nuclei) still formed – more later

6. Behavior of elements depend on the number and configuration of electrons

7. Distribution of elements in whole earth depends on solar system formation (loss of volatiles like H, He, noble gases, C, O) and redistribution of nonvolatile elements by density during differentiation of the Earth into core/mantle and crust.

8. Earth’s biogeochemical cycles depend fundamentally on the composition of the Earth – what elements it is made up of and the lithosphere cycles that operate to distribute them between earth and the earth’s surface/ocean/atmosphere system where we live We will ignore (for the most part) these long-term (multi-millennial) cycles that exchange material between inner and outer earth components (mantle/surface) ; these cycles are driven by lithosphere processes with energy from radiodecay of nuclides in the mantle/core

9. Why we need biogeochemistry: Elements take different forms in different parts of the Earth System so transfers from one sphere to another involve change of chemical form or change of phase Liquid/ dissolved ion Solid/liquid/ dissolved ion Solid Gas/liquid

10. The Bio-part: Biogeocemical Cycles involve elements essential to life: C, N, O, P, S, H, and SiThese are interrelated through stoichiometry of living things, e.g. Ocean Redfield Ratio (Redfield, Ketchum, and Richards,1963)106CO2+ 16 NO3- + HPO42-+ 122 H2O + 18 H+ <=> (CH2O)106(NH3)16(H3PO4) + 138 O2Ocean C:N:P is ~ 106:16:1 (not always!!)On land, ratios are different – C:N and C:P can be much higher because plant structural material is built out of C-rich materials like cellulose (C6H12O2)Stoichiometry means: Element balance (same number on each side of the equation – reactants and products)Charge balance (electroneutrality) Charge balance

11. Example: equation of State for Gases – needed to understand expression of gas abundance in the atmosphere and relate volume occupied to amount and state variables (physical conditions:T,P) Thermodynamics (relationship between energy and chemistry): PV = nRT Partial pressures (how we express gas concentrations) pT = (p1 +p2 +…pn)V = (n1 + n2 +…Nn)RT e.g. Carbon dioxide = 360 ppm or 360 x 10-6 moles CO2 per mole air or 360 matmospheres CO2 per 1 atmosphere total air pressure

13. Units – mixing ratio ppm = parts per million Or 1 mole CO2 per mole air Covariance of O2 and CO2 is fixed by stoichiometry of biological (photosynthesis) and nonbiological (dissolution) and human (fossil fuel burning) processes

14. Dynamic Equilibrium The reaction coefficient (Keq) - relates the concentrations (activities) of reactants and products for a system at dynamic equilibrium. Can be determined from energetics of the system: ΔG = ΔG0 + RT ln (Keq) The reaction quotient - ratios of actual concentrations; tells you if you are at equilibrium or not. Equilibrium is dynamic – reactions are occurring but rates balance so there is no macroscopic change in the system observable; the distribution of reactants and products represents maximizes entropy in the universe

15. CH4 + 2O2 CO2 + 2H2O Current concentrations of gases in the atmosphere: CH4: 2 ppm (2 x 10-6 mol CH4/mol air) O2 : 21% (0.21 mol O2/mol air) CO2 380 ppm (380 x 10-6 mol CO2/mol air H2O: ~1% water vapor in atmsophere (.01 mol H2O/mol air) Where a is activity = concentration or ideal pressure times a correction factor to represent the role of molecular interactions that might interfere with the ability to participate in reactions –’nonideal’ conditions. In this course, we will ignore these complications and use concentration or partial pressure rather than activity. Keq is determined (from thermodynamic constants – see notes) to be ~10140 Is methane in the atmosphere at equilibrium??? i.e. Q equals Keq?

16. Lovelock and Gaia • The atmosphere is not at chemical equilibrium with ocean/land surfaces • Gases are present in nonequilibrium concentrations because they are continuously produced and consumed by biological processes • Feedbacks between those gases and the life that produces them maintain the atmosphere in a state conducive to the continuation of life

17. CHEMISTRY IS NOT ENOUGH BY ITSELFCharacteristic times for exchange – often more important than chemical kinetics; can be used to constrain chemical kinetics. We need to understand the physics of atmospheric and ocean transport to explain observed distribution of chemicals in the environment Stratosphere years Interhemispheric transport 1 year Troposphere month hours Planetary Boundary Layer Land slow transfers (water) Surface Mixed layer hours years Deep Ocean (interbasin transport 100- 1000 yrs) (from Henning Rodhe Chapter 4in Earth System Science)

18. The Global Atmospheric Circulation Timescales for mixing of gases in the atmosphere: Zonal – weeks Equator to pole – weeks Across the equator – 1 year (because of convergence at the ITCZ) Barbara Summey, NASA Goddard Visualization Lab

19. Tools to Study Complex BGC Systems (1) Identify the Elements of the system and how they interact (Biogeochemical budget) (2) Determine the characteristic response times for chemical and physical transformations (how fast do the elements interact, and how fast will a change affect the system? how fast are interactions compared to physical mixing constraints in earth reservoirs? (3) Identify possible feedback loops/interactions with other biogeochemical cycles or climate conditions - will they tend to amplify (positive feedback) or damp (negative feedback) changes to the system? For very complex systems with many interacting elements, we need to construct computer models to predict how the system will respond to a disturbance

20. Global water cycle – change of phase only (H2O, gas, liquid, solid)

21. Reservoir = the amount of the material of interest in a given form. A reservoir has a finite capacity and in studying it we define its exchanges with other reservoirs. Water in a lake, water in the atmosphere, water in the ocean. Example: Water in the atmosphere: 13 x 103 km3 Water in the ocean: 1.37x 109 km3 • Flux = the amount of material added to (Source) or removed from (Sink) the reservoir in a given period of time. An example of a flux is the evaporation of water from the surface of a lake or ocean (in which the ocean reservoir is a source of water for the atmosphere). Ocean fluxes: In: rivers + precipitation = (40,000 + 385,000) x 103 km3 a-1 Out: evaporation = 425,000 ) x 103 km3 a-1 Atmosphere fluxes: (out)= precipitation = (111,000 + 385,000) x 103 km3 a-1 Residence Time – how long to empty the reservoir (Reservoir size/Flux out) Ocean: 37,000 year • A reservoir that is at Steady State shows no change in mass with time (Sources or fluxes in = Sinks or fluxes out)

22. Rodhe’s three key terms for expressing the dynamics of cycling for geochemical reservoirs • Turnover time This is the time it would take to empty (or fill) the reservoir. At steady state this is the amount of material in reservoir divided by the sum of all fluxes out or sum of all fluxes in. • Mean (average) Residence Time This is the average time spent in the reservoir by individual atoms - measurable as the age of atoms leaving the reservoir • Mean Age This is the average the average time spent in the reservoir by all the atoms currently in the reservoir (measurable as the age of atoms in the reservoir) For a HOMOGENEOUS reservoir at STEADY STATE (i.e. not changing in amount with time) all three of these terms are equal. However, we make our biggest mistakes in defining systems to be homogeneous (i.e. all elements in the reservoir have the same probability of leaving) when they are not.

23. Flux = 200 g/year Flux =50 g/year 0.05 kg Flux = 50 g/year Flux = 150 g/year 5 kg 4.950 kg Flux = 200 g/year Flux = 150 g/year Case 1. single homogeneous pool at steady state (often assumed but often not really what is there). Case 2. two components each with different turnover times; same overall inventory and total flux in/flux out System must be defined carefully! – calculate the overall turnover time, mean residence time and mean age for these two cases

24. Heat Capacity  and The Earth's Heat Storage

25. Response Time depends on heat content (heat capacity*Temperature) This is why the climate system has lags – e.g we say that there is still warming built in to the climate system.

26. How can you change climate? • Change incoming solar radiation • (solar constant) • 2. Change amount of absorbed solar radiation (albedo) • Change amount of absorbed and re-radiated longwave radiation (greenhouse gases) • Change distribution of absorbed energy (internal climate system) Incoming Sunlight heats Earth Changes can be amplified or dampened by feedbacks within the climate system

27. A major negative feedback controlling Earth’s temperature Increase incoming energy Planet will warm Warmer planet will lose more heat Planet temperature will stabilize Negative feedback Incoming Sunlight heats Earth Planet Temperature Rate of Heat Loss

28. A positive feedback example – ice and clouds and the amount of sunlight reflected from earth Reflected light Incoming sunlight Lighter surface (ice, cloud) reflects more sunlight (absorbs less) Cooler Dark surface absorbs more sunlight Warmer

29. This is a feedback because the amount of ice depends on the temperature of the planet – colder = more ice Amount of sunlight reflected Temperature of planet Amount of ice Step 1 Identify the elements of the system that interact. In this case the output (reflected sunlight) affects the input (sunlight absorbed at planet surface), both influence planet’s temperature

30. If more sunlight is reflected, temperature cools Amount of sunlight reflected Temperature of planet Amount of ice This is a negative coupling – an increase in reflected sunlight causes a decrease in the temperature of the planet

31. If temperature cools, there is more ice formed Amount of sunlight reflected Temperature of planet Amount of ice This is a negative coupling – a decrease in temperature causes an increase in the amount of the planet’s surface covered by ice

32. If there is more ice, there will be a greater amount of sunlight reflected Amount of sunlight reflected Temperature of planet Amount of ice This is a positive coupling – an increase in ice coverage causes an increase in the amount incoming sunlight reflected

33. Overall this is a POSITIVE feedback;it has an even number of negative couplings Amount of sunlight reflected Temperature of planet (+) Amount of ice Positive feedbacks are destabilizing – they amplify changes –an increase in the amount of sunlight reflected will end up increasing the amount of reflected sunlight; temperatures will keep dropping and ice cover will increase

34. The “Snowball Earth” hypothesis – At one point (~700 million years ago) Earth may have been entirely ice covered and much colder because of a ‘runaway’ ice – reflectance feedback Which begs the question – how did we get out of the frozen state? There must be another process at work …more later on

35. Example – Feedbacks between climate and Greenhouse Gases (more gases = warmer average temperature) Natural greenhouse gas processes (CO2, CH4, …) Atmospheric Stock of GHG (radiative forcing) • What is the magnitude of these GHG feedbacks? • How do they affect climate predictions? Water (vapor, clouds, ice) Global Warming

36. Ice core data imply large, positive greenhouse gas feedbacks Ice cores effectively integrate over all feedback mechanisms and record the net result. CH4 Emily Stone (NSF) Temperature changes recorded in ice cores are larger than the forcing we can understand that operates on the same timescales/periodicity (orbital variations). Greenhouse gases change after temperature in ice cores, implies positive feedback Temp CO2 0 600,000 Age (yr B P) EPICA Ice Core 650,000 year record Torn and Harte. Missing feedbacks, asymmetric uncertainties, and the underestimation of future warming, GRL, 2006.

37. Temperature DTi DTF atm CO2 How is the feedback effect quantified? g g= feedback gain, g < 1 Total g = sum of independent g’s = gCO2 + gCH4 + gGCM

38. CO2 (ppm) Temperature anomaly (ºC) www.nersc.gov CO2 (ppm) T (ºC) kya The gain gfor greenhouse gases can be calculated from GCMs and Ice Cores GCM Output If we doubled CO2 and ran the climate model to equilibrium what would be the temperature change? Ice Core Data Torn & Harte, GRL, 2006

39. Summary: Studying Global Biogeochemical Cycles and the underlying biogeochemical processes that control them, gives us constraints for understanding how those cycles can be altered – by humans, by climate variation, etc. Tools for analyzing Global Biogeochemical Cycles – define reservoirs appropriately - determine if the system can be approximated by steady state - determine the turnover time -various methods for quantifying fluxes (ocean, land, etc) - how do those fluxes couple to climate/other biogeochemical cycles? - what are the major feedbacks/couplings and are they positive/negative? - can we quantify feedbacks via calculation of gain?

40. Earth System Science is different • The Earth has history and it influences what we observe today • We have only one (no replication, except in models) • Things do not often change in isolation Paleoclimate (sunlight variation effects) (effects of volcanoes)

41. Tracers– another useful tool Isotopes (or tracer analogs) of an element provide very useful information for global biogeochemical cycles Same chemistry, slightly different rates of reaction How do isotopes of H and O help us understand the water cycle?

42. Isotopes of an element have the same number of protons (therefore chemistry) but different numbers of neutrons (mass) 12C 98.9% (6 protons, 6 neutrons) 13C 1.1 % (6 protons, 7 neutrons) 14C ~1.1x 10-10 % (6 protons, 8 neutrons) Isotopes that are unstable decay radioactively – 14C decays to 14N Isotopes of carbon contain different information: 13C variations – patterns in the environment reflect mass-dependent fractionation (partitioning among phases at equilibrium and differences in reaction rates), mixing of sources 14C variations – Reflects time since isolation from exchange with atmosphere or mixing of sources

43. http://www.science.uottawa.ca/~eih/ch1/ch1.htm#enai

44. Isotope Fractionation • Fractionation refers to the processes that cause partial separation of two isotopes of the same element, producing reservoirs with different ratios of rare to abundant isotopes. • Mass dependent fractionation • equilibrium isotope fractionation - due to differences in bond energies of isotopes in compounds • kinetic isotope fractionation- due to differences in average velocity or reaction rates of different isotopes • Both depend only on the mass of the isotope and are called mass dependent fractionation; both will fractionate, say 18O/16O about twice as much as 17O/16O because the mass difference between 18 and 16 is ~ twice that of 17 and 16 • Mass-independent fractionation – generally requires higher energies – e.g. those associated with photolysis. At higher temperatures/energies, the distribution of isotopes among bond types is random rather than by bond energy, so that there is no difference between 18O/16O and 17O/16O

45. Why do isotopes fractionate? Small differences in mass mean small differences in chemical bond strength and vibrational frequency (n) Vibrations are quantized, E(vib) = hn(n+1/2) n=0,1,2,3 ½(hn) (n=0) is called the Zero-Point Energy (ZPE) Potential energy

46. Property H2O D2O Density (20 C) 0.9982 1.1050 Temperature of greatest density 4.0 11.6 Mole volume (20C) cm3/mole 18.049 18.124 Melting point (760 torr, C) 0.00 3.82 Boiling point (760 torr, C) 100.00 101.42 Vapor pressure (at 100C in torr) 760.00 721.60 Viscosity (@20.2 C in centipoises) 1.00 1.26 Ionic product a 20C 1x10-14 0.16 x 10-14 Example: Characteristic constants of H2O and D2O (From Hoefs, 1973, Stable isotope geochemistry)

48. [H216O]vapor + [H218O]liquid ↔ [H216O]vapor + [H218O]liquid The equilibrium constant for this reaction is In this case Keq = the fractionation factor, a More generally, a = K (1/n), where n is the maximum number of exchange positions that are equivalent (for example n =2 for H in H2O).

50. What accounts for the differences in stable isotope content among reservoirs?