What Is Carbon 14 Used For

Carbon fourteen

Electrical Accuse

George B. Arfken , ... Joseph Priest , in University Physics, 1984

Example ii Carbon 14 Beta Decay

Carbon fourteen, ¹⁴ ₆C, has six protons in its nucleus and is formed in our atmosphere by cosmic ray bombardment of nitrogen 14, ^xiv ₇Due north. Carbon 14 is unstable and undergoes radioactive decay past emitting a beta particle (an electron) and an antineutrino (zero mass, aught charge) and becoming nitrogen (seven protons):

${}_{\mathrm{six}}^{\mathrm{xiv}}\text{C}\to {}_{-\mathrm{one}}e+{}_{\mathrm{seven}}^{14}\text{Northward}\text{+}\overline{v}$

In this process a neutral particle in the ^l4C nucleus, chosen a neutron, is transformed into three particles—a positively charged proton, a negatively charged electron, and a neutral antineutrino. The proton, the electron, and the antineutrino are created from the neutron in the reaction. Merely although both positive and negative charges are created, the cyberspace charge (+ half dozen before = - one + 7 after = + 6 later on) remains the same. Again the subscripts represent exact conservation of electric accuse.

In Example 2 we saw the creation of positive and negative electric charge. Now let'south consider the reverse of charge creation: charge annihilation.

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Radiometric Dating

Marvin Lanphere , in Encyclopedia of Concrete Scientific discipline and Applied science (Tertiary Edition), 2002

III.F The Radiocarbon Method

Carbon-14 is a short-lived radioactive nuclide compared to the longer-lived radioactive isotopes described previously (Table I). But, ¹⁴C (radioactive carbon or radiocarbon) dating is and then important to archaeology, anthropology, geology, and other fields that the method merits discussion.¹⁴C is continually produced in the upper atmosphere by interactions of cosmic ray neutrons with ¹⁴N. The cosmic ray neutrons interact with the stable isotopes of nitrogen, oxygen, and carbon, simply the interaction with stable ¹⁴Due north is the most important of these reactions. The reaction of a neutron with ^fourteenDue north produces ¹⁴C and a proton is emitted from the nucleus. The atoms of ¹⁴C are oxidized inside a few hours to ^fourteenCO, which has an atmospheric lifetime of several months. The ^xivCO is in turn oxidized to ¹⁴CO₂. These molecules of CO_ii take a relatively long atmospheric lifetime of ∼100 years which allows the ^xivCO_two to exist well mixed and achieve a steady-land equilibrium in the atmosphere. This equilibrium is maintained by product of ¹⁴C in the temper and continuous radioactive disuse of ^fourteenC. Molecules of ¹⁴CO₂ enter institute tissue as a result of photosynthesis or by absorption through the roots. The rapid cycling of carbon betwixt the temper and biosphere allows plants to maintain a ¹⁴C activity approximately equal to the activity of the atmosphere. However, the isotopes of carbon are fractionated by physical and chemic reactions that occur in nature. This fractionation introduces small systematic errors in radiocarbon dates. In addition to ¹⁴C, carbon includes ii stable isotopes, ¹²C and ¹³C. The mass-dependent fractionation can be eliminated past measuring the ¹²C/¹⁴C ratio on a mass spectrometer. Animals that feed on the plants as well acquire a abiding level of radioactivity due to ¹⁴C. When the found or creature dies, the assimilation of ^xivC from the atmosphere stops and the activity of ¹⁴C decreases due to radioactive decay. ¹⁴C undergoes β-decay to ¹⁴Northward with a half-life of 5730 years.

At some time afterward the death of an organism the activity of ^fourteenC in dead tissue can be compared with the activity of ^fourteenC in shortly living tissue to yield a carbon-fourteen or radiocarbon date for the sample. Notation that the radiocarbon method is completely different from the accumulation clocks described above. Those clocks are based on the accumulation of radiogenic daughter isotope produced by radioactive decay of a parent isotope. The radiocarbon method is based on the corporeality of ¹⁴C remaining after radioactive decay of ¹⁴C. The radioactivity of carbon extracted from plant or beast tissue that died t years ago is given by:

(17) $\text{A}={\text{A}}_0{\text{e}}^{-\lambda \text{t}},$

where A is the measured activity of ¹⁴C in the sample in units of disintegrations per minute per gram of carbon, and A₀ is the activity of ¹⁴C in the same sample at the time the establish or animal was alive. The carbon-14 age of a sample containing carbon that is no longer in equilibrium with ¹⁴C in the atmosphere or hydrosphere is obtained by solving equation 17 for t:

(18) ${\text{log}}_{\text{e}}\left(\text{A}/{\text{A}}_0\right)=-\lambda \text{t}\text{or}\text{t}=\mathrm{ane}/\lambda {\text{log}}_{\text{e}}\left({\text{A}}_0/\text{A}\right).$

The carbon-14 dating method depends on special assumptions regarding A₀ and A. These are (1) that the rate of ¹⁴C production in the upper temper is constant and has been contained of time, and (2) that the rate of assimilation of ¹⁴C into living organisms is rapid relative to the rate of decay. As volition be shown below the first supposition is not easily satisfied. It is known that the neutron flux increases with altitude above the globe's surface and that the flux is near four times greater in polar areas than at the equator. However, the ¹⁴C activeness is known to be independent of latitude.

The first ¹⁴C ages were measured in the 1940s on elemental carbon in amorphous class ("carbon black"). Nonetheless, the product of massive amounts of artificial radiocarbon from atmospheric testing of nuclear weapons in the 1950s complicated the apply of elemental carbon for depression-level ¹⁴C measurements. Methods were subsequently developed to count the decay of ¹⁴C chemically converted to purified CO_two, hydrocarbon gases and liquids. Modern laboratories can measure ¹⁴C ages on organic thing as old every bit 40,000 to 50,000 years. A few laboratories accept developed the adequacy to measure ages every bit old every bit seventy,000 years on larger samples. In the 1970s the advent of accelerator mass spectrometry resulted in a major boost in detection efficiency. The amount of carbon required for a measurement was reduced from grams to milligrams, and the counting fourth dimension was reduced from days or weeks to minutes. Information technology was idea that the increased detection sensitivity might extend the maximum historic period datable by the radiocarbon method from the routine forty,000 to l,000 years to peradventure 100,000 years. However, it turned out that contamination past younger or modern carbon, which unremarkably is introduced past chemical or biological activity subsequent to the death of the sample and also during sample preparation limits accelerator mass spectrometry ages to the same twoscore,000 to fifty,000 years. It has not been possible to appointment with high precision samples less than about 300 years old except under special weather. The natural fluctuations in ¹⁴C production combined with the release of large quantities of fossil fuel CO_ii and production of "flop" ¹⁴C from atmospheric nuclear testing accept made the measurement of young radiocarbon ages very hard.

In the 1950s discrepancies between radiocarbon ages and truthful ages were noted. A long-term trend with superimposed shorter-term deviations in the ¹⁴C time calibration indicated that the supposition of constant product charge per unit of ^xivC in the atmosphere probably was not truthful in detail. The corporeality of offset betwixt ¹⁴C ages and calendar ages has been calibrated for the past 11,800 years by measuring ¹⁴C ages on wood for trees for which true ages could be determined by counting yearly growth rings. Past convention ages are referred to AD 1950   which is equivalent to 0 years BP (earlier present). From 11,800 to 24,000 years BP radiocarbon ages have been calibrated against uranium–thorium disequilibrium ages of corals or varve-counted marine sediments. The discrepancy between uncorrected ¹⁴C years and calendar years at 24,000 years is iii,700 years. Figurer programs are available to calculate the starting time betwixt ¹⁴C and calendar years.

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FUNDAMENTALS OF NUCLEAR MAGNETIC RESONANCE

FRANK A. BOVEY , PETER A. MIRAU , in NMR of Polymers, 1996

one.5.2 CARBON-xiii Chemical SHIFTS

Carbon-xiii chemical shifts are very sensitive to molecular structure and geometry, quite apart from the influence of substituent groups. This is peculiarly clearly revealed in the ¹³C spectra of paraffinic hydrocarbons, a course to which many important polymers belong. In contrast to paraffinic protons, which embrace a range of merely almost 2 ppm (Fig. 1.nineteen), carbon chemic shifts are spread over more then 40 ppm (Fig. 1.20). This makes ¹³C spectroscopy a powerful means for the report of such materials. The empirical ordering of such effects (theoretical agreement lags far behind) may be done in terms of α, β, and γ effects (27,28) (δ and ε furnishings are further refinements which we omit here). Table i.2 shows a series of data for simple hydrocarbons that illustrate the α effect. Here, nosotros are to observe °C carbon and ask what happens on adding carbons α to this one. We see a regular deshielding of nearly ix ± i ppm for each added carbon, except in neopentane, where crowding apparently reduces the outcome.

TABLE 1.2. The α Effect in ¹³C Shieldings

Structure	δC (ppm)	α Effect (ppm)
°CH3 − H	−i.2	—
°CH3 − CH3	five.9	8.0
°CHii − (CH3)2	16.i	10.2
°CH − (CHiii)3	25.2	nine.1
°C − (CH3)4	27.9	2.vii

In Table 1.three we see examples of the outcome of carbons added β to the observed 1, °C. The event is of similar magnitude to that produced by α carbons, a particularly difficult bespeak theoretically. For nonterminal carbons the outcome is similar, but reduced in magnitude if °C is a branch point.

TABLE 1.three. The β Effect in ¹³C Shieldings

Structure	δC (ppm)	β Effect (ppm)
°CH3 − αCH3	5.9	—
°CH3 − αCH2 − βCHthree	fifteen.6	9.7
°CH3 − αCH−(βCHiii)ii	24.3	8.seven
°CHiii − αC−(βCH3)iii	31.5	7.2
αCH3 − °CH2 − αCH3	16.1	—
αCHiii − °CH2 − αCH2 − βCH3	25.0	8.ix
αCH3 − 0CH2 − αCH−(βCH3)2	31.8	6.8
αCH3−°CH2 − αC−(βCH3)iii	36.7	4.ix
(αCH3)two−°CHii −αCH3	25.2	−
(αCHiii)2−°CH2 −αCHtwo−βCHiii	29.nine	four.7
(αCHthree)ii− °CH2 −αCH-(βCH3)2	34.1	4.ii
(αCH3)2−°CH2 −αC-(βCHthree)3	38.one	4.0

Finally, we must consider the γ effect (Table one.4), which, although smaller than α and β effects, is of particular interest because dissimilar them it is a shielding effect and too is clearly dependent on molecular conformation (29). The γ effect is found to exist produced by elements other than carbon which are three bonds removed from the observed carbon and to show a correlation with the electronegativity of this element (vide infra).

Table 1.iv. The γ Effect in ¹³C Shieldings

Structure	δC (ppm)	γ Effect (ppm)
°CHiii −αCHtwo−βCH3	15.half-dozen	—
°CH3 − αCH2 − βCHtwo − γCH3	xiii.two	−2.4
°CH3 − αCH2 − βCH-(γCH3)two	xi.3	−1.9
°CH3 − αCH2 − βC-(γCHthree)3	8.8	−2.five
αCH3 − °CH2 − αCH2 − βCHiii	25.0	—
αCH3 − °CHii − αCHii − βCH2 − γCH3	22.six	−2.4
αCH3−°CH2 -°CH2−βCH-(γCHthree)2	20.seven	−i.9
αCH3−°CH2 −αCH2−βC-(γCH3)three	18.8	−i.9

For a better understanding of the γ effect, consider the staggered conformers of three of the compounds in Table 1.4, represented in Scheme one. In the example of butane, we split up the observed average γ shift by the gauche conformer content (ca. 0.45) and find a shielding value of −v.iii ppm per gauche interaction. Information technology is proposed that this shielding occurs when the four-carbon three-bond system is in or nearly a gauche state. We shall see in afterwards word that when the dihedral bending decreases from the gauche value of 60° toward the eclipse conformation (0°), the shielding increases to most 8 ppm, while every bit the angle increases the effect decreases, becoming zero beyond about 90°–100°. Shielding differences of this magnitude tin can exist readily observed and measured in the solid state and provide a welcome substantiation of crystal structures derived from X-ray diffraction.

We shall see later that the gauche shielding value of −five.3 ppm explains quite accurately, among other observations, the dependence of carbon chemic shifts on stereochemical configuration in many chiral compounds and macromolecules. For the more crowded compounds in Scheme 1, 2-methylbutane and 2,two-dimethylbutane, the effect appears somewhat smaller.

The rules and regularities embodied in Tables 1.two, 1.3, and one.4 and in the γ-gauche issue enable one to predict as well equally interpret carbon chemical shifts, and thereby to exam proposed structures and conformations. A more detailed discussion of the manner of making such prediction is given elsewhere (30) and will not business organization us further here.

Anterior effects are of bully importance in carbon-13 chemic shifts of non-paraffinic compounds and macromolecules. The transmission of anterior effects is illustrated by the schematic spectra in Fig. 1.25. Theory (31,32) suggests that the deshielding influence of electronegative elements or groups should be propagated down the concatenation with alternating effect, decreasing as the 3rd power of the distance. This prediction is non borne out. It is true that the shielding of C₁ shows the expected dependence on the electronegativity of the fastened atom X. Note particularly the effect of F and O. The relative effects of the halogens are also what one expects. Note that iodine actually causes C_one to be more shielded than C₅. We notice, however, that the shielding of C₂ is independent of electronegativity. The substitution of any element for the H of pentane produces about the same deshielding event at C₂, i.e., 8 to 10 ppm. Such commutation causes a shielding of C_iii by ca. 2-6 ppm, an effect we have already noted and discussed when the substituting element is carbon merely which is caused by other elements also. In fact, carbon is at the low end of this two to vi-ppm range; N and O produce a larger "γ" effect.

Effigy i.25. Schematic carbon-xiii spectra illustrating the transmission of anterior furnishings in direct-concatenation aliphatic compounds. All spectra represented as proton decoupled so that carbon resonances appear as singlets.

In Fig. 1.xx nosotros saw that olefinic and aromatic carbons are strongly deshielded compared to those of saturated carbon chains—an effect parallel to that for protons (Fig. one.xix) simply much larger. Theoretical explanations (31,32) involve the paramagnetic screening term rather than diamagnetic shielding or ring electric current considerations (33–35). Carbonyl carbon atoms testify fifty-fifty greater deshielding (Fig. i.20).

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The Oceans and Marine Geochemistry

J. Lynch-Stieglitz , in Treatise on Geochemistry, 2003

half-dozen.16.5.1 Radiocarbon

Carbon-14 (radiocarbon) is produced in the temper and decays with a half-life of 5,730 yr. Natural radiocarbon levels in the sea are highest in the surface body of water, where they partially equilibrate with atmosphere. Deep-h2o concentrations are highest in newly formed NADW and decrease to the everyman values in the deep Pacific, where the deep waters accept longest been out of contact with the atmosphere (Stuiver et al., 1983). Benthic foraminifera record the radiocarbon concentrations of the deep water in which they grow, only the ¹⁴C in the foraminifera will decay over time. In order to reconstruct the radiocarbon content at the time the foraminifera were living, radiocarbon measurements tin be fabricated on benthic and planktonic foraminifera from the same interval in the cadre (Broecker et al., 1988, 1990; Shackleton et al., 1988) in gild to determine the age of the benthic foraminifera. Withal, the radiocarbon content of the planktonic foraminifera will reflect non only the age of the planktonic foraminifera, but the initial radiocarbon content of the near-surface water in which they grew. The initial radiocarbon in surface waters and in near-surface waters can exist significantly less than expected for equilibrium with the atmosphere due to the relatively long fourth dimension required for isotopic equilibrium between carbon in the surface ocean and temper. I must also business relationship for the fact that the radiocarbon content in the atmosphere changes with time, and will as well touch the initial radiocarbon content of surface waters (Adkins and Boyle, 1997).

Besides complicating the use of concurrent benthic and planktonic radiocarbon concentrations to assess water mass historic period is the requirement that the benthic and planktonic foraminifera measured at the same interval actually be the same age. If the peak abundance of planktonic and benthic foraminifera occurs at different depths in the cadre, bioturbation can innovate a spread in the planktonic–benthic ages in cases with low accumulation rate, which is unrelated to changes in lesser-h2o historic period (Peng et al., 1977; Broecker et al., 1999). While these problems tin can be overcome by using cores with high sedimentation rate (e.g., Keigwin and Schlegel, 2002), the widespread use of this techniques has been limited primarily by the identification of suitable cores with loftier enough foraminiferal abundances for accurate radiocarbon determinations. Sediments with high aggregating rates tend to either accept low foraminiferal abundances (due to dilution by terrigenous textile), or be in high-productivity regions of the sea, where surface waters are expected to have variable and relatively low initial radiocarbon content.

Some other mode to measure the deep-water radiocarbon ventilation age in the past is past using deep-dwelling house benthic corals (Adkins et al., 1998; Mangini et al., 1998; Goldstein et al., 2001). Here, the age of the coral can be independently determined using uranium-series dating allowing the radiocarbon content of the coral to exist used to determine the radiocarbon content of deep water at the time the coral grew. This method is unaffected by bioturbation, but is currently limited by the availability of deep-sea benthic corals.

Changes in the ventilation of the body of water can likewise be assessed by looking at how the ¹⁴C content of the atmosphere has inverse through time (e.thou., Hughen et al., 1998). The atmospheric ^xivC content is controlled both by the rate of ^fourteenC production in the atmosphere and the charge per unit at which ^xivC is transferred into the ocean. While a subtract in the ventilation of the ocean causes a buildup of radiocarbon in the atmosphere, one must also account for changes in production earlier interpreting the atmospheric radiocarbon changes in terms of ocean circulation (Marchal et al., 2001).

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Surface and Ground H2o, Weathering, and Soils

F.Thousand. Phillips , Chiliad.C. Castro , in Treatise on Geochemistry, 2003

5.fifteen.4.2.7 Carbon-14

Carbon-14 is produced in the temper past a depression-energy cosmic-ray neutron reaction with nitrogen. It decays dorsum to ¹⁴Due north with a half-life of v,730 year. The production charge per unit is ∼2×10⁴  cantlet   m^−two  s⁻¹, the highest of all cosmogenic radionuclides. The rate is so high, because nitrogen is the most arable element in the atmosphere and also has a very big thermal neutron assimilation cross-section. Radiocarbon activity in the temper is the result of complicated exchanges betwixt terrestrial reservoirs (primarily vegetation), the bounding main, and the atmosphere. The combination of varying cosmic-ray fluxes (due to both varying solar and terrestrial magnetic field modulations) and varying substitution betwixt reservoirs has caused the atmospheric activeness of radiocarbon to fluctuate past approximately a factor of 2 over the past 3×10^iv  yr (Bard, 1998). The current specific activity is 0.23   Bq (g C)⁻¹, rendering radiocarbon easily measurable by gas proportional or liquid scintillation counting, or by accelerator mass spectrometry (AMS). Radiocarbon measurements are ordinarily reported as "percent modern carbon," indicating the sample specific action as a percentage of the 0.23 Bq (gC)^−one modern atmospheric specific activeness. In add-on to the natural cosmogenic product of ¹⁴C, the isotope was besides released in big amounts past atmospheric nuclear-weapons testing in the 1950s and 1960s (Figure 2(b)).

Natural radiocarbon was kickoff detected past Libby in the mid-1940s (Arnold and Libby, 1949; Libby, 1946), merely the first applications to subsurface hydrology were not attempted for another decade (Hanshaw et al., 1965; Münnich, 1957; Pearson, 1966). These early on investigators discovered that radiocarbon shows clear and systematic decreases with catamenia altitude that can be attributed to radiodecay, but as well exhibits the effects of carbonate mineral dissolution and precipitation reactions. Quantification of residence fourth dimension is not possible without correction for additions of nonatmospheric carbon. Numerous approaches to this problem were attempted, including uncomplicated empirical measurements (Vogel, 1970), simplified chemical equilibrium mass residue (Tamers, 1975), mass balance using δ¹³C as an analogue for ^fourteenC (Ingerson and Pearson, 1964), and combined chemical/δ^thirteenC mass balance (Fontes and Garnier, 1979; Mook, 1980).

The correction methods cited above were mainly intended to business relationship for carbonate reactions in the vadose zone during recharge, although in practice they take often been applied to reactions along the groundwater flow path also. It is at present the full general consensus that the preferred arroyo to treating the continuing reactions of dissolved inorganic carbon during menstruum is geochemical mass transfer/equilibrium models (Kalin, 2000; Zhu and Potato, 2000). The most ordinarily employed model is NETPATH (run into Capacity 5.02 and 5.fourteen; Plummer et al., 1991). This uses a backward-calculated solute mass balance, constrained by uncomplicated equilibrium considerations, to calculate mass transfers of solutes from stage to stage between two sample points.

The vast majority of groundwater radiocarbon investigations accept extracted the dissolved inorganic carbon from the water for measurement. However, dissolved organic carbon (DOC) has been sampled in a limited number of studies (Potato et al., 1989; Purdy et al., 1992; Tullborg and Gustafsson, 1999). In principle, sampling DOC could avoid the circuitous geochemistry of inorganic carbon. In reality, Doc consists of a big number of organic species from various sources (with different original ¹⁴C activities), of varying chemical stability, and of varying reactivity with the solid phase. Routine use of Doc for groundwater age tracing may 1 mean solar day show advantages over inorganic carbon, but this will require considerable work separating and identifying the appropriate component of the Medico for this application.

Radiocarbon is an indispensable tool in the age tracing of groundwater. Both its ease of use and its half-life arrive the method of choice for many investigations. However, results must always exist approached with caution. The reliability of results depends strongly on the complication of the geochemistry. In some cases groundwater may show footling evidence of chemical development afterward recharge and measured radiocarbon activities can be accustomed at face up value. In other cases, the isotopic composition of carbon may be strongly altered by numerous surface reactions that are difficult to quantify, and interpretations may be speculative, at best. Interpretations must be evaluated on a case-by-case basis.

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Surface and Ground Water, Weathering, and Soils

R. Amundson , in Treatise on Geochemistry, 2003

five.01.v.ane.3 Processes and isotope composition of pedogenic carbonate formation

In arid and semi-arid regions of the globe, where precipitation is exceeded by potential evapotranspiration, soils are incompletely leached and CaCO₃ accumulates in significant quantities. Figure 18 illustrates the global distribution of carbonate in the upper meter of soils. As the figure illustrates, there is a sharp boundary between calcareous and noncalcareous soil in the USA at about the 100th summit. This long-recognized boundary reflects the soil water rest. Jenny and Leonard (1939) examined the depth to the top of the carbonate-bearing layer in soils by establishing a climosequence (precipitation gradient) forth an east to due west transect of the Great Plains (Effigy xix) . They observed that at constant hateful average temperature (MAT) below 100 cm of mean average precipitation (MAP), carbonate appeared in the soils, and the depth to the height of the carbonate layer decreased with decreasing precipitation. An analysis has been made of the depth to carbonate versus precipitation relation for the entire USA (Royer, 1999). She institute that, in full general, the relation exists broadly but as the control on other variables betwixt sites (temperature, soil texture, etc.) is relaxed, the strength of the human relationship declines greatly.

Figure 18. Global distribution of pedogenic carbonate (source http://www.nrcs.usda.gov/technical/worldsoils/mapindx/).

Figure xix. (a) Schematic view of plant and soil profile changes on an due east (correct) to west (left) transect of the Great Plains and (b) measured depth to elevation of pedogenic carbonate forth the aforementioned gradient (source Jenny, 1941).

In add-on to the depth versus climate tendency, there is a predictable and repeatable tendency of carbonate amount and morphology with time (Gile et al. (1966); Figure 20) due to the progressive aggregating of carbonate over time, and the ultimate infilling of soil porosity with carbonate cement, which restricts further downward movement of water and carbonate.

Figure 20. Soil carbonate morphology and amount versus time for: (a) gravelly and (b) fine-grained soils (Gile et al., 1966) (reproduced by permission of Williams and Wilkins from Soil Sci. 1966, 101, 347–360).

The controls underlying the depth and amount of soil carbonate swivel on the water balance, Ca⁺ⁱⁱ availability, soil CO₂ partial pressures, etc. Arkley (1963) was the first to characterize these processes mathematically. His work has been greatly expanded past McFadden and Tinsley (1985), Marian et al. (1985), and others to include numerical models. Figure 21 illustrates the general concepts of McFadden and Tinsley'southward numerical model, and Figure 22 illustrates the results of model predictions for a hot, semi-arid soil (see the figure heading for model parameter values). These predictions more often than not mimic observations of carbonate distribution in desert soils, indicating that many of the key processes have been identified.

Effigy 21. Diagram of compartment model for the numerical simulation of calcic soil evolution. Compartments on left represent solid phases and compartments on right correspond aqueous phases. Line A represents atmospheric precipitation, line B represents dust influx, line C represents the transfer of components due to dissolution and atmospheric precipitation, line D represents transfer betwixt aqueous phases, line Due east represents downward movement of solutes due to gravitational flow of soil h2o, line F represents evapotranspirational water loss, and line G represents leaching losses of solutes (McFadden et al., 1991) (reproduced by permission of Soil Scientific discipline Order of America from Occurrence, Characteristics, and Genesis of Carbonate, Gypsum and Silica Accumulations in Soils, 1991).

Figure 22. Predicted pattern of Holocene pedogenic carbonate aggregating in a cmⁱⁱ column in a semi-barren, thermic climate (leaching alphabetize=three.5 cm). External carbonate flux rate=one.v×ten⁻⁴  g   cm⁻²  year⁻¹, p _CO2 =i.5×10^{−three.three}  atm in compartment 1 increasing to 10^−2.5atm in compartment 5 (20–25 cm). Below compartment v, the p _CO2 decreases to a minimum value of x⁻⁴atm in compartment 20 (95–100 cm). Dotted line shows carbonate distribution at t=0. Gray area indicates terminal faux distribution. Depth^*=absolute infiltration depth in <2 mm fraction (McFadden et al., 1991) (reproduced by permission of Soil Science Club of America from Occurrence, Characteristics, and Genesis of Carbonate, Gypsum and Silica Accumulations in Soils, 1991).

The full general equation describing the formation of carbonate in soils is illustrated by the reaction

${\text{CO}}_{\mathrm{ii}}+{\text{H}}_2\text{O}+{\text{Ca}}^{+2}={\text{CaCO}}_{\mathrm{three}}+\mathrm{ii}{\text{H}}^{+}$

From an isotopic perspective, in unsaturated soils, soil CO₂ represents an infinite reservoir of carbon and soil water an infinite reservoir of oxygen, and the δ^xiiiC and δ¹⁸O values of the pedogenic carbonate (regardless of whether its calcium is derived from silicate weathering, atmospheric sources, or limestone) are entirely fix past the isotopic composition of soil CO_ii and H₂O. Here we focus mainly on the carbon isotopes. However, briefly for completeness, we outline the oxygen-isotope processes in soils. The source of soil H₂O is atmospheric precipitation, whose oxygen-isotope composition is controlled by a complex set of physical processes (Hendricks et al., 2000), but which commonly shows a positive correlation with MAT (Rozanski et al., 1993). Once this h2o enters the soil, information technology is subject to transpirational (largely unfractionating) and evaporative (highly fractionating) losses. Barnes and Allison (1983) presented an evaporative soil water model that consists of processes for an: (i) upper, vapor transport zone and (2) a liquid water zone with an upper evaporating front. The model describes the circuitous variations observed in soil water versus depth following periods of extensive evaporation (Barnes and Allison, 1983; Stern et al., 1999). Pedogenic carbonate that forms in soils generally mirrors these soil water patterns (e.yard., Cerling and Quade, 1993). In general, in all but hyperarid, poorly vegetated sites (where the evaporation/transpiration ratio is high), soil CaCO₃ δ¹⁸O values roughly reflect those of precipitation (Amundson et al., 1996).

The carbon-isotope model and its variants take, it is fair to say, revolutionized the apply of soils and paleosols (see Chapter 5.xviii) in paleobotany and climatology. Some of the major achievements and uses of the model include the post-obit.

•

The model adequately describes the observed increases in the δ¹³C value of both soil CO₂ and CaCO_three with depth (Figure 17).

•

The model clearly provides a mechanistic understanding of why soil CO₂ is enriched in ¹³C relative to plant inputs (steady-state diffusional enrichment of ¹³C).

•

The model indicates that for reasonable rates of CO_two production in soils, the δ¹³C value of soil CO_two should, at a depth within 100 cm of the surface, represent the δ^xiiiC value of the standing biomass plus 4.four‰. The δ¹³C of CaCO_iii will likewise reflect this value, plus an equilibrium fractionation of ∼10‰ (depending on temperature). Therefore, if paleosols are sampled below the "atmospheric mixing zone," whose thickness depends on CO₂ product rates (Figure 17), the δ¹³C value of the carbonate will provide a guide to the by vegetation (Cerling et al., 1989).

Cerling (1991) recognized that Equation (fourteen), if rearranged and solved for C _atm, could provide a means of utilizing paleosol carbonates for reconstructing past atmospheric CO_two partial pressures. To practice and so, the values of the other variables (including the δ¹³C value of the atmospheric CO₂—meet Jahren et al. (2001)) must be known, which for soils of the distant past is not necessarily a fiddling problem. Nonetheless, an active research field has developed using this method, and a compilation of calculated atmospheric CO₂ levels is emerging (Ekart et al.,1999; Mora et al., 1996), with estimates that correlate well with model calculations by Berner (1992).

Cerling's (1984) approach to modeling stable carbon isotopes in soil CO₂ has been expanded and adapted to other isotopes in soil CO_ii: (i) ^xivCO₂—for pedogenic carbonate dating (Wang et al., 1994; Amundson et al., 1994a,b) and soil carbon turnover studies (Wang et al., 2000) and (ii) C¹⁸O¹⁶O—for hydrological tracer applications and, more than importantly, every bit a means to constrain the controls on global atmospheric CO₂–¹⁸O budgets (Hesterberg and Siegenthaler, 1991; Amundson et al., 1998; Stern et al., 1999, 2001; Tans, 1998). The processes decision-making the isotopes, and the complication of the models, greatly increment from ¹⁴C to ^eighteenO (run across Amundson et al. (1998) for a detailed account of all soil CO₂ isotopic models).

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Sediments, Diagenesis, and Sedimentary Rocks

A. Lerman , N. Clauer , in Treatise on Geochemistry, 2007

7.16.vi.2 Long-Term Trends of Carbonate Sediments

The δ^xiiiC values of the Precambrian limestones and dolomites, representing a menstruum of more than 3,000   Ma, are shown in Figure 15. Shields and Veizer (2002) point out that the ages of some of the data points are not well constrained on a resolution timescale of fifty–100   Ma, some may correspond local marine environments that are not feature of a world average at their fourth dimension, and there may exist differences between the shallower-water platform sediments and those of the pelagic body of water. These uncertainties probably business relationship for the big spread of the δ¹³C values in the Paleoproterozoic and Neoproterozoic, around the fourth dimension of the 2 major cold periods or glaciations in the Earth's early history. The Paleoarchean values of δ^xiiiC≈0‰ of calcites and dolomites are consistent with the occurrence of biological fractionation of carbon since 3,800–three,900   Ma. Because at the Earth's surface temperatures (5–25   °C), the ¹³C/¹²C fractionation betwixt calcite and the HCO_three ⁻ ion is less than i‰ (Salomons and Mook, 1986), the δ^thirteenC of marine carbonates is interpreted as that of ocean water. A general trend for calcites and dolomites is shown in Figure xvi where the individual data points were averaged by successive 100   Ma time intervals. Relatively large standard deviations (±1σ) of some of the means are due to the large spreads effectually those ways, as axiomatic in the private data plotted in Figure fifteen. It is noteworthy that the two minerals evidence little difference in their δ¹³C values through the long fourth dimension of the Precambrian, as mentioned earlier: the hateful difference δ^xiiiC_dol−δ¹³C_dol≈0.4‰, simply its extremes are +5‰ at 2,050   Ma and about −iii‰ at 650   Ma and at ii,700–2,750   Ma.

Figure 15. δ¹³C values of Precambrian dolomites (four,914 data points) and calcites (iv,266 points), from Shields and Veizer (2002). Plot from data courtesy of J. Veizer (Academy of Ottawa).

Figure sixteen. Mean values of δ¹³C of Precambrian dolomites and calcites from Figure 15 past 100   Ma time intervals (thick line), ±1 standard deviation (sparse lines). Annotation the differences betwixt the vertical scales in Figures xv and 16. Plot from information courtesy of J. Veizer.

The δ¹³C of drapery carbon is approximately −5‰ to −7‰ and in the absenteeism of life and its photosynthetic capabilities, this would also be the isotopic composition of seawater. A higher value is a result of the biological fractionation between organic thing and CaCO₃, removal of each into the sediments, and subsequent subduction of the bounding main-floor sediments into the mantle. If the carbon input to the sea in the Paleoarchean was similar to the drapery carbon δ¹³C, and then the organic thing fraction of the sedimentary carbon would have been about 17%, assuming the biological fractionation ε=30‰, equally discussed before. The increase in the δ¹³C of the carbonates in the time interval from nearly 2,400 to 2,200   Ma is usually interpreted as an indication of greater biological production and storage of organic thing in sediments and, consequently, an increase in the oxygen concentration in the atmosphere. An increment in δ¹³C from 0‰ to ten‰ in the carbonates would represent to an increase in the stored organic carbon fraction to nearly l%, assuming the aforementioned values of the input and biological fractionation.

The carbon isotope record of Phanerozoic time, since the beginning of the Cambrian Flow, is shown in Figure 17. Although the current age of the base of the Cambrian is placed at 542   Ma (Gradstein et al., 2004), its age guess has varied from 470 to 610   Ma since 1937 (Harland et al., 1990). It is conceivable that the large spread of δ^thirteenC values near the Precambrian–Cambrian boundary reflects poor time constraints for some of the data, as pointed out by Shields and Veizer (2002). A number of long-term trends tin can be distinguished in the Phanerozoic:

Figure 17. δ^xiiiC values of Phanerozoic carbonates. Cambrian to Cretaceous (black • and blue +) information from Veizer et al. (1999; courtesy of J. Veizer); Jurassic to Contempo (red ×) from Katz et al. (2005). VPDB scale. Stratigraphic timescale from Gradstein et al. (2004). Thick line: consecutive 10-Ma-interval means (5   Ma in the most recent 10   Ma interval); thin lines are ±1 standard deviation.

•

In the Cambrian, an increase in the δ¹³C of carbonates during a period of about 60   Ma, coincides with the expansion of the multicellular organisms that may be related to a greater biological production at the lower levels of the food chain.

•

Although a slight subtract in the Ordovician occurs before than the Late Ordovician glaciation, in that location is a secular increase in the δ¹³C during the side by side 180   Ma from the Ordovician to the Carboniferous–Permian boundary that broadly coincides with the emergence of college plants and massive storage of organic carbon as coal in the Carboniferous.

•

The next 100   Ma include glaciations of Gondwana near and higher up the Carboniferous–Permian purlieus and a major extinction at the Permian–Triassic transition. This time represents a slightly declining trend, with much scatter, of the carbonate δ^xiiiC. Still, some detailed δ¹³C profiles across the Permian–Triassic boundary have provided internally consistent trends and data for the modeling of changes in atmospheric CO_two and O₂ at that time (Berner, 2006).

•

From the Jurassic to the Miocene, a period from almost 200 to twenty   Ma, in that location is a slight increment in the δ¹³C, although it is smaller than an increase in the Paleozoic over a similar length of fourth dimension. Within the Jurassic to the Miocene data, in that location are periods of elevated δ^xiiiC values that lasted about 5–10   Ma and shorter periods of ≤i   Ma that are associated with oceanic anoxic events (Katz et al., 2005).

•

The decrease in the δ^thirteenC of carbonates at to the lowest degree during the last thirty   Ma to the present and an increase in the δ¹³C of organic carbon during the same period (Figures 14 and 17) correspond a trend that differs from the past and it translates into a smaller biological fractionation between the carbonate and organic carbon in oceanic sediments. It has been variably proposed that a greater weathering of former organic matter has increased the delivery of isotopically lighter inorganic carbon forming from the remineralization of ¹³C-depleted organic matter and information technology has been, consequently, the crusade of the decreasing δ¹³C values of marine carbonates. However, the δ¹³C values of country plants exercise non back up this interpretation, every bit discussed in an earlier section.

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Isotope Ratio Studies Using Mass Spectrometry

Michael E. Wieser , Willi A. Brand , in Encyclopedia of Spectroscopy and Spectrometry, 1999

Standards

The standard for carbon isotope abundance measurements is based on a Cretaceous belemnite sample from the Peedee formation in Due south Carolina, USA. The original fabric is no longer available. It has been replaced by the convention that NBS xix, a carbonate material, has a value of + one.95% versus PDB. This new scale is termed Five-PDB (Vienna-PDB). The IAEA distributes a number of secondary standards including graphite (USGS24) with a δ^thirteenC value of − 15.99% Five-PDB, oil (NBS-22) at − 29.74% Five-PDB, and calcium carbonate (NBS-18) with a value of − 5.01% 5-PDB.

The oxygen isotope composition of carbonates is also ordinarily referenced to the ^eighteenO/¹⁶O-isotope ratio of V-PDB. Information technology is not used as a reference for δ¹⁸O analyses of igneous or metamorphic rocks. The departure between V-SMOW and V-PDB δ¹⁸O values is approximately 30% (δ¹⁸O_V-SMOW = 1.03091 × δ¹⁸O_Five-PDB + thirty.91).

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A VERSATILE COMPUTER ORIENTED LIQUID SCINTILLATION COUNTING Organization USING THE DOUBLE RATIO TECHNIQUE

D.Due south. Drinking glass , in Organic Scintillators and Scintillation Counting, 1971

Counting Weather

Discriminator and gain settings for carbon-14, tritium and the external standard are shown in Table ii. All standards are counted under the aforementioned conditions. The ratios of interest are shown in Table 3. The relationship between efficiency and the corresponding ratios for carbon-14 and tritium in organic scintillator are shown in the Figures 1–4. The graphs show that although a very polish correlation is obtained between the sample channels ratio for both tritium and carbon-14, a more intricate curve is obtained for the respective AES ratio/efficiency graphs. Even with a large number of standards it has been found hard to obtain a good fit with a polynomial regression assay for these curves.

Tabular array 2. Instrument Settings

		Isotope	Gain (%)	Window (v)
Channel	1	Carbon-xiv	5.7	l-1000
Aqueduct	two	Tritium	50	l-thousand
Channel	3	External Standard	2.0	350-1000

Tabular array 3. Calibration Ratios

Isotope	Sample Channels Ratio	External Standard Channels Ratio
Carbon-14	Cyberspace Sample Counts (C2)	Cyberspace Standard Counts (C1)
	Internet Sample Counts (C1)	Net Standard Counts (C3)
Tritium	Net Sample Counts (C1)	Net Standard Counts (C3)
	Net Sample Counts (C2)	Net Standard Counts (C1)

Fig. 1. Carbon-xiv sample channels ratio against efficiency

Fig. 2. Carbon-fourteen external standard ratio against efficiency

Fig. 3. Tritium sample channels ratio confronting efficiency

Fig. 4. Tritium external standard ratio against efficiency

If the ratios are inverted, equally we reported previously, (2) a satisfactory fit is obtained with a single quintic polynomial over the whole range. Even so, as a consequence of increased computer orientation of the arrangement, information technology was decided to fit the AES curves using a series of quadratic regression equations. This becomes feasible only when data handling can be reduced to an accented minimum by writing and applying suitable figurer programs.

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Sediments, Diagenesis, and Sedimentary Rocks

B.B. Sageman , T.West. Lyons , in Treatise on Geochemistry, 2003

vii.06.3.3.2 Stable carbon isotopes of OM (δ¹³C)

Despite the wide multifariousness of controls on the carbon isotope composition of preserved OM and biogenic CaCO_three (east.yard., Hayes, 1993; Kump and Arthur, 1999), it is possible under certain circumstances to argue that a few variables are ascendant and thus to relate changes in δ¹³C_org to trends in paleoproduction. Every bit argued originally by Scholle and Arthur (1980), Lewan (1986), Arthur et al. (1988), and others, changes in the δ^thirteenC of preserved marine algal OM and biogenic CaCO_iii reflect changes in the isotopic composition of dissolved inorganic carbon in surface waters, which on short to intermediate timescales (<1   Myr) may exist controlled by the residue betwixt net respiration and net burial of OM in sediments. For example, positive shifts in the δ¹³C of organic carbon dominantly sourced from marine photoautotrophs accept been interpreted to reverberate elevated burial fluxes of OM related to global increment in primary productivity (Arthur et al., 1987, 1988), whereas negative shifts have been interpreted to reflect recycling of respired CO₂ in a more localized reservoir (e.chiliad., Saelen et al., 1998; Murphy et al., 2000a; Rohl et al., 2001). For more detailed recent reviews of the controls on C-isotope fractionation see Kump and Arthur (1999), Hayes et al. (1999), and papers in Valley and Cole (2001).

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What Is Carbon 14 Used For,

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