A Methane-Based Time Scale for Vostok Ice William F

Quaternary Science Reviews 22 (2003) 141–155

A methane-based time scale for Vostok ice William F. Ruddimana,*, Maureen E. Raymob a Department of Environmental Sciences, University of Virginia, Clark Hall, Charlottesville, VA22903, USA b Department of Earth Sciences, Boston University, Boston, MA 02215, USA Received 7 February 2002; accepted 21 May 2002

Abstract

Tuning the Vostok methane signal to mid-July 301N insolation yields a new ice-core gas time scale. This exercise has two rationales: (1) evidence supporting Kutzbach’s theory that low-latitude summer insolation in the northern hemisphere controls the strength oftropical monsoons, and (2) interhemispheric CH 4 gradients showing that the main control oforbital-scale CH 4 variations is tropical (monsoonal) sources. The immediate basis for tuning CH4 to mid-July insolation is the coincident timing ofthe most recent (pre-anthropogenic) CH4 maximum at 11,000–10,500 calendar years ago and the most recent July 301N insolation maximum (all ages in this paper are in calendar years unless speciﬁed as 14C years). The resulting CH4 gas time scale diverges by as much as 15,000 years from the GT4 gas time scale (Petit et al., Nature 399 (1999) 18 429) prior to 250,000 years ago, but it matches fairly closely a time scale derived by tuning ice-core d Oatm to a lagged insolation 18 signal (Shackleton, Science 289 (2000) 1897). Most offsets between the CH4 and d Oatm time scales can be explained by assuming 18 that tropical monsoons and ice sheets alternate in controlling the phase ofthe d Oatm signal. 18 The CH4 time scale provides an estimate ofthe timing ofthe Vostok CO 2 signal against SPECMAP marine d O, often used as an 18 index ofglobal ice volume. On the CH 4 time scale, all CO2 responses are highly coherent with SPECMAP d O at the orbital 18 periods. CO2 leads d O by 5000 years at 100,000 years (eccentricity), but the two signals are nearly in-phase at 41,000 years (obliquity) and 23,000 years (precession). The actual phasing between CO2 and ice volume is difﬁcult to infer because of likely SST 18 overprints on the SPECMAP d O signal. CO2 could lead, or be in phase with, ice volume, but is unlikely to lag behind the ice response. r 2002 Elsevier Science Ltd. All rights reserved.

0. Introduction between ice-core signals and insolation forcing, as well as other climatic responses such as ice volume. Accurate Ice cores contain an incredible variety ofclimatic phasing (lead/lag) information is vital to efforts to signals that have had a major impact on climate establish cause-and-effect (forcing-and-response) cli- research. Over time spans of400,000 years or more, matic relationships. ice-core signals show obvious cyclic behavior concen- Section 1 ofthis paper reviews the history of trated within the range ofastronomically calculated development oftime scales forVostok ice record, changes in Earth’s orbit, from 20,000 to 100,000 years. including estimates ofthe uncertainties associated with To date, however, it has proven difficult to realize the each method. Section 2 sets out the rationale for full promise of these ice-core signals because the time creating a new Vostok gas time scale based on monsoon scale has not been sufficiently accurate. Without a forcing of atmospheric methane (CH4) concentrations. consistently high level ofaccuracy (to within a few Section 3 makes an initial assessment ofthis new CH 4 thousand years), it is difficult to assess how each ice-core time scale and concludes that it is preferable to the GT4 signal is partitioned among the orbital periods. More (gas) time scale ofPetit et al. (1999). Section 4 focuseson critically, it is impossible to fix the orbital-scale phasing differences between the CH4 time scale and the generally 18 similar d Oatm time scale ofShackleton (2000) and *Corresponding author. Fax: +1-804-982-2137. concludes that the CH4 time scale provides a more E-mail addresses: [email protected] (W.F. Ruddiman), plausible explanation for the small offsets. Section 5 uses [email protected] (M.E. Raymo). the CH4 time scale as a basis to define the orbital-scale

18 phasing ofCO 2 relative to SPECMAP d O. Finally, ibility record in Southern Ocean core RC11-120, based Section 6 summarizes factors that make it difficult to on the assumption that the marine magnetic suscept- infer the phasing of CO2 against ice volume. ibility signal at 411S is a proxy for dust deposition. Pichon et al. (1992) and Waelbroeck et al. (1995) correlated deuterium (dD) in Vostok ice with diatom- 1. History of time scales for Vostok ice based estimates ofsea surface temperature in marine cores from middle latitudes of the Southern Ocean based Early time scales for the Vostok ice record across the on the assumption that mid-latitude changes in sea- last full interglacial–glacial cycle were derived from ice- surface temperature correlate with air-temperature flow/ice-accumulation models; these are based on the changes at 801S. 18 initial rate ofaccumulation ofsnow in the area feeding Sowers et al. (1993) tuned the d O ofatmospheric O 2 ice to the Vostok core site, and on subsequent ice in Vostok ice to a marine d18O target signal in Pacific thinning determined from an ice-flow model (Lorius core V19–30. First, they tried to remove the imprint of et al., 1985). The initial rate ofaccumulation is bottom-water temperature changes from the benthic 18 18 18 estimated from d O values in the ice using modern foraminiferal d O signal to isolate d Osw (the mean spatial relationships. The flow model also requires d18O value ofseawater) as an ice-volume tuning target. assumptions about the form of the surface and bedrock Raymo and Horowitz (1996) used an inferred marine 13 topography upstream from the ice-core site. d C proxy for past CO2 values to correlate with the This initial time scale for Vostok ice raised concerns Vostok CO2 record, but also checked matches ofVostok 18 18 because the next-to-last deglaciation occurred more than d Oatm against marine d O values. Shackleton (2000) 18 10,000 years before the seemingly equivalent feature in tuned Vostok d Oatm to a synthetic orbital insolation the marine d18O signal and was thus out ofphase with signal used as a target. He relied on the age ofthe major 18 respect to the presumed insolation forcing. This d Oatm change at the last deglaciation in annually mismatch was lessened when Jouzel et al. (1993) altered layered GISP ice to set the phase ofthe precession and 18 the time scale ofLorius et al. (1985) by incorporating obliquity components ofthe d Oatm signal relative to considerably higher accumulation rates in features insolation. considered equivalent to marine isotopic stages 6 and 5. Each technique for estimating ages carries inherent Petit et al. (1999) published signals from the full uncertainties. These errors are difficult to establish for length ofthe Vostok ice core. For the upper 110,000 ice-flow/ice-accumulation models, but were estimated by years ofthe record, they used a slightly adjusted version Petit et al. (1999) as ranging from 75000 to 15,000 years ofthe ice-flow/ice-accumulation time scale fromJouzel in different parts of the Vostok record. Uncertainties in et al. (1993). They pinned the rest ofthe ice time scale counting annual layers can be estimated by comparing (termed ‘‘GT4’’) to levels assumed to represent 110,000 different signals in the same core or the same signal in and 423,000 years ago, based on inferred correlations to nearby cores (Meese et al., 1997). These range from 100 the marine d18O record. Between these points, they to 200 years back to 12,000 years ago to about 5% ofthe estimated ages by calculating ice thinning based on a absolute age down to 55,000 years ago (the level by flow model, with adjustments for basal melting and which annual layering cannot be reliably detected). sliding. A second method for dating ice cores is by Techniques that link ice-core signals to marine records counting annual layers preserved in rapidly deposited produce errors in the correlation process, with progres- ice. This method cannot be used in the slowly deposited sively larger errors for successively more dissimilar Vostok ice, but time scales developed in the last 15,000 signals. These techniques also implicitly incorporate years from Greenland ice can be transferred to Vostok the 3000-year to 5000-year uncertainty in orbital tuning using globally correlative climate signals such as ofthe marine record (Imbrie et al., 1984). Correlations 18 18 methane. of d Oatm to marine d Osw also introduce an un- A third approach is to correlate (‘‘tune’’) climate- certainty of 7500 years in the turnover time ofO 2 in the sensitive proxies within Vostok ice to climatic signals in atmosphere (Sowers et al., 1993). In addition, correla- marine sediments. In effect, this method transfers the tions between ice and gas phases within Vostok ice may orbitally tuned time scale ofthe marine record to the ice be in error by 72000 years (Sowers et al., 1993). core. Marine records are tuned by matching the filtered 41,000-year and 23,000-year components ofmarine d18O signals to assumed forcing at the corresponding orbital 2. Rationale for tuning ice-core methane (CH4) to orbital periods, with lags appropriate to the long time constants insolation ofice sheets. Several such correlations to the marine record have We present here a new gas time scale for Vostok based been attempted. Petit et al. (1999) correlated the dust on a simple and well-justified approach: tuning the concentration in Vostok ice with a magnetic suscept- methane (CH4) gas signal in Vostok ice to mid-July W.F. Ruddiman, M.E. Raymo / Quaternary Science Reviews 22 (2003) 141–155 143 insolation at 301N. This method is based on two 800 assumptions. The first assumption is that the CH4 signal is mainly a response to the strength oflow-latitude

monsoons, with secondary input from boreal sources July 30 (Raynaud et al., 1988; Chappellaz et al., 1990; Blunier 700 et al., 1995; Brook et al., 1996). Strong summer 500 ° monsoons deliver moisture to southeast Asia, India, N Insolation (W/m and northern Africa, and the monsoonal rains ﬁll lakes and wetlands that become sources ofmethane as (ppbv) vegetation decays in reducing conditions. Modeling of 4 600

CH Y. D. CH4 gradients between Greenland and Antarctic ice 490

over the last 40,000 years indicates that the tropics have 2 ) been the dominant source ofmethane over orbital time Insolation scales (Chappellaz et al., 1997; Brook et al., 2000). 500 CH4 The second assumption is that tropical monsoons are driven by low-latitude summer insolation with a 480 dominant orbital precession signal. This assumption 5.000 10.000 15.000 builds on the orbital monsoon hypothesis first proposed Calender Years BP by Kutzbach (1981). This hypothesis has since been supported by a wide array ofevidence, including Fig. 1. Comparison ofCH 4 trends between 15,000 and 5000 years ago tropical lake levels (Street-Perrott and Harrison, 1984; (based on a composite trend from the GRIP ice core assembled by Chappellaz et al., 1997) and July insolation at 301N (from Berger and Kutzbach and Street-Perrott, 1985), Nile-induced stra- Loutre, 1992). YD marks the Younger Dryas event. tification ofthe Mediterranean Sea (Rossignol-Strick et al., 1982), and influxes offreshwater diatoms blown from African lakes into Atlantic sediments (Pokras and 21 summer solstice is often used as a reference for Mix, 1987). analyzing insolation changes, but mid-July insolation This tuning exercise requires two key choices ofan appears to provide the optimal match to the GRIP CH4 insolation ‘‘tuning target’’: (1) a critical latitude, and (2) peak (Fig. 1). GCM analyses ofsummer monsoon a critical season or month. To guide this choice, we forcing often choose the June/July/August season, compare the timing ofinsolation signals calculated in centered on mid-July, as the critical season (Prell and calendric years against the calendric age ofthe most Kutzbach, 1987). recent CH4 maximum in the GRIP ice core from We chose 301N as the critical insolation latitude to Greenland as determined from counting annual layers. balance the combined effect of several land masses The most recent CH4 maximum in the GRIP ice core heated by low-latitude insolation. North Africa lies prior to the anthropogenic era is a broad maximum almost entirely in the tropics (o231N), but Asia extends from 14,500 to 8500 years ago (Fig. 1). This peak is through the subtropics and far into higher latitudes. interrupted by an abrupt CH4 minimum during the Across much ofthe north tropics and subtropics, a Younger Dryas event near 12,000 years ago, and is different choice of latitude would not significantly alter followed by a smaller minimum just before 8000 years the shape (or timing) ofthe tuning target. But at ago. Despite these interruptions, the wave-like form of latitudes above 501N, insolation signals begin to take on the orbital-scale peak remains evident. Peak CH4 values more ofthe signature ofthe 41,000-year obliquity signal, are centered near 11,000 to 10,500 years ago. Attempts and this can nudge specific precessional maxima and to refine this estimate ofthe age ofthe CH 4 maximum minima toward slightly younger or older ages. using a spline fit would be hampered by the presence of This rationale for tuning methane to July 301N the abrupt Younger Dryas event and the rapid deglacial insolation might be challenged in several ways. One CH4 increase near 14,500 years ago. At this level in the basis for challenge is that some climatic proxies linked to ice, the age offset between the CH4 gas and the ice in tropical monsoons reach maximum responses later than which it is trapped is about 50 years (Blunier et al., the ice-core CH4 maximum. Radiocarbon dating shows 1995). that closed-basin lakes in North Africa reached their 14 This age for the CH4 maximum matches the timing of highest levels between 9000 and 6000 C years ago the most recent mid-July insolation maximum at low (Street-Perrott and Harrison, 1984) and that Nile River and middle latitudes (Berger and Loutre, 1992). Because runoff from monsoonal regions of northeast Africa a family of precession curves exists for each season and created organic-rich sapropel muds in the Mediterra- month due to the slow progression ofthe solstices and nean Sea from 9000 to 7000 14C years ago (Thunell and equinoxes around the elliptical orbit, it is necessary to Williams, 1989). Even when converted to calendar years, select one ofthese monthly signals as critical. The June these responses occurred 10,000 to 7500 years ago, well 144 W.F. Ruddiman, M.E. Raymo / Quaternary Science Reviews 22 (2003) 141–155 after the July insolation maximum from 11,000 to sediments at high latitudes or in shallow marine 10,500 years ago. Similar lags occur in upwelling proxies sediments at any latitude, could be another possible and pollen influxes off the Somali coast of eastern Africa source ofatmospheric CH 4. The rapid changes in CH4 (Prell, 1984; Prell and van Campo, 1986). These lags, during the last 15,000 years shown in Fig. 1 could be which occur in and around Africa, may reflect external interpreted as lending credibility to this idea. To date, climatic influences on the African monsoon such as the however, ice-core studies have not favored this possibi- lingering effects of northern ice sheets or North Atlantic lity. The character ofthe record in Fig. 1 is more easily SST (deMenocal and Rind, 1996), or the delayed timing interpreted as one in which cold-climate CH4 minima ofthe 23,000-year SST maximum in the eastern tropical interrupt longer-term orbital-scale CH4 maxima, rather Atlantic from which the monsoon draws latent heat than as warm-climate pulses ofexcess CH 4 expelled (McIntyre et al., 1989). from clathrates. In addition, Brook et al. (2000) noted On close inspection, however, many ofthe best-dated that even the most rapid CH4 declines were spread over African lake records resemble the methane trends 200 to 300 years, an interval significantly longer than the plotted in Fig. 1, showing major increases in monsoonal abrupt shifts of other ice-core signals towards colder moisture 14,500 years ago, maximum lake levels by conditions, and consistent with the time scale for 11,000 years ago, and high lake levels until 5000 to 4000 ecosystem adjustments to suddenly imposed reductions years ago (Gasse, 2000). In contrast, the Mediterranean in tropical moisture. sapropels occur well toward the end ofthis moist In any case, the slowly deposited ice at Vostok acts as interval, suggesting that they are a lagging indicator of a low-pass filter and smoothes most millennial-scale monsoon conditions. CH4 fluctuations, leaving behind a record mainly of In any case, the recent CH4 maximum in the GRIP ice orbital-scale changes. Over this time scale, Chappellaz core used to set the timing ofour tuning exercise is well et al. (1997) inferred that any clathrate contribution to dated, and other evidence from Asia supports this ice-core CH4 must have arrived as a relatively steady timing. Planktonic foraminiferal d18O signals in South input over orbital-scale intervals. As such, this input China Sea cores show negative departures from typical would be indistinguishable from monsoonal inputs. d18O values during the last deglaciation (Wang et al., Our method carries less uncertainty than any other 1999). These deviations are interpreted as salinity ice-dating technique. The CH4 signal plotted in Fig. 1 is decreases caused by excess river inflow caused by south in error by at most 150750 years (Meese et al., 1997). Asian monsoons. This inferred monsoon signal first Visual inspection shows that it would be difficult to shift appeared 15,000 years ago, markedly increased in our choice ofJuly as the critical insolation month by strength just before 13,000 years BP, and gradually more than 7300 to 400 years without harming the ended after 9000 years ago. Given the large wetland match shown in Fig. 1. (Such a shift equates to less than expanse in southern Asia, Asian CH4 sources probably a week ifconverted to equivalent time in the preces- dominated the ice-core CH4 signal. sional cycle.) In addition, monthly (and seasonal) A second possible challenge to our rationale is that insolation signals have no significant dating error across part ofthe CH 4 signal comes from boreal wetland the time range examined here (Berger and Loutre, 1992). sources far from regions controlled by tropical mon- As a result, errors in our choice oftuning target are soons. But CH4 gradients between Greenland and unlikely to exceed 7500 years. Antarctic ice suggest that tropical methane sources were The largest potential source oferror in our procedure twice as large as those north of30 1N during both the (assuming the validity ofour basic rationale) arises from last-glacial CH4 minimum and the late-deglacial CH4 the process oflinking maxima and minima in the CH 4 maximum (Chappellaz et al., 1997; Brook et al., 2000). signal with the precisely dated insolation signal (Table 1). This indicates that tropical monsoons dominated the The Vostok CH4 signal is sampled at an average interval CH4 signal through a range ofclimatic changes. of o1500 years over the last 350,000 years, with the In addition, the mid-July insolation peak that spacing ranging from o1000 years in the upper part of produces maximum amounts ofCH 4 from tropical the record to about 2000 years in the older portion. This monsoons at the 23,000-year cycle should also drive a sample spacing introduces an average uncertainty of coincident CH4 influx from boreal regions. Mid-summer 7750 years across the record, with a range from 7500 heating ofthe Asian continent should warm the chilly years in the upper part to 71000 years in the older part. northern wetlands, stimulate a briefmid-summer pulse Combined with the 7500-year error in our initial choice ofmicrobial activity, and liberate maximum amounts of ofthe tuning target, we estimate that the errors boreal CH4. The mid-July phasing ofthe CH 4 maximum associated with our method could optimally be as small in Fig. 1 is consistent with the idea ofcoincident mid- as 71000–2000 years. summer CH4 influxes from tropical and boreal sources. Across some intervals, however, our method could A third potential criticism is that melting ofCH 4- have larger errors. Precession-driven CH4 maxima and bearing clathrates, either those frozen in terrestrial minima are not well developed during the glacial parts W.F. Ruddiman, M.E. Raymo / Quaternary Science Reviews 22 (2003) 141–155 145

Table 1 tion uncertainties occur within low-amplitude CH4 CH4 time scale for Vostok ice oscillations during full-glacial conditions. In addition, 3 Depth in ice (m) Age (10 yr) Feature correlated some CH4 peaks are interrupted by briefminima, but the basic precessional wave pattern remains apparent 149.2 0 — 321.2 11 Max through most ofthe record. We call this time scale the 506.4 22 Min ‘‘CH4 time scale’’ (Table 1). 664.1 32 Max The Vostok record below 3180 m ice depth (350,000 748.3 45 Min years age) is omitted because neither the CH4 signal nor 860.6 58 Max the 301N insolation tuning target has sufficiently 1087.2 70 Min 1237.2 82 Max distinctive amplitude variations within this age range 1338.2 93 Min to permit credible tuning. It is also unclear whether the 1526.2 104 Max Vostok record reaches the bottom ofmarine isotopic 1557.4 115 Min stage 11 (Petit et al., 1999). 1893.4 131 Max Cross-spectral analysis confirms that the strong 2030.9 137 Midpoint 2137.1 149 Max 23,000-year power in the tuned CH4 signal is in phase, 2177.3 161 Min and highly coherent, with precessional insolation 2363.0 174 Max (Fig. 2b–d). The largest mismatches in amplitude occur 2437.0 185 Min on CH4 minima, most ofwhich are rectified at minimum 2525.0 196 Max values of400–420 ppb, even when insolation fallsto very 2557.7 208 Min 2621.7 219 Max low values. Amplitude mismatches between insolation 2698.0 230 Min and CH4 also occur near terminations, with somewhat 2771.2 240 Max deeper minima just prior to deglaciations and dispro- 2830.4 251 Min portionately strong CH4 maxima early in the following 2857.5 265 Max interglaciations. These mismatches suggest that factors 2911.4 278 Min 2944.5 290 Max other than insolation also effect the amplitude of early 2994.5 301 Min interglacial CH4 maxima. 3051.5 312 Max The main difference between the two signals is the 3054.5 320 Min concentration ofCH 4 power at 100,000 and 41,000 3078.5 328 Midpoint years, periods that are not prominent in the July 301N 3109.5 334 Max 3138.5 338 Midpoint insolation tuning signal (Fig. 2b). The 100,000-year 3165.5 344 Min signal is in phase with eccentricity, largely as a result ofthe common eccentricity modulation ofboth CH Feature correlated specifies maxima, minima or transition midpoints in 4 and ofthe precessional insolation signal used as the both the 301N insolation curve and the Vostok CH4 record used as control points for correlation, with linear interpolation of intervening tuning target. The 41,000-year CH4 signal lags well levels. behind insolation at the obliquity period (Fig. 2c) and is nearly in phase with the 41,000-year component of SPECMAP d18O. This phasing suggests that the 41,000- ofthe Vostok record, such as the intervals around year CH4 signal largely reflects an overprint ofice 40,000, 150,000, and 260,000 years ago. During these volume (or some other glacial boundary condition) on intervals, it is difficult to link the CH4 record to methane emissions from tropical and/or boreal wetlands precessional insolation. As a result, errors larger than (Prell and Kutzbach, 1987). 5000 years (1/4 ofa precessional cycle) are not The methane record plotted at the CH4 time scale is impossible. For the rest ofthe record, errors greater compared with that ofthe GT4 time scale ofPetit et al. than 2000 years are unlikely. Finally, as is the case for all (1999) in Fig. 3. Because the CH4 time scale used CH4 tuning techniques, our method is vulnerable to the gas for tuning, we use the GT4 gas time scale for this assumption that the phase ofthe monsoon may have comparison, rather than the GT4 time scale for ice. The varied through time because ofvariable influences of GT4 gas time scale was derived by subtracting 2000 to other parts ofthe climate system (ice sheets, SST, etc). 6000 years from the time scale for ice, with larger differences during glacial intervals of slow ice accumulation (Petit et al. 1999). 3. Initial assessment of the CH4 time scale For most ofthe record back to 250,000 years ago, the CH4 and GT4 time scales are similar, with age A time-series comparison shows the tuned correlation departures ofno more than 5000 to 10,000 years, and between Vostok CH4 and 301N July insolation (Fig. 2a). generally only a few thousand years (Fig. 3a). The ages Both signals show distinct amplitude modulation at the ofmany CH 4 maxima and minima in the GT4 time scale precessional cycle. As noted above, the major correla- only needed to be shifted by 1000 to 3000 years to create 146 W.F. Ruddiman, M.E. Raymo / Quaternary Science Reviews 22 (2003) 141–155

800 July 30N Vostok methane (ppbv) )

2 Vostok CH 520 4

600

480 400 July insolation (W/m 440 200 (a) 0 50 100 150 200 250 300 350 kyr 41 kyr 100 23 kyr 5 104 2 105

July 30N 4 Vostok CH 4 10 4 1.5 105

3 104 BW 5 power 1 10

2 104 Rel.

5 104 1 104

0 (b) 1.0 0 10

0.8 95% CL

0.6

0.4 Coherency

0.2

Phase (deg.) -90

0.00 0.05 0.10 (d) Frequency kyr-1

Fig. 2. Comparison ofVostok CH 4 (plotted on the CH4 time scale) and July 301N insolation for the last 350,000 years: (a) Time series; (b–d) Cross- spectral analysis; (b) power spectra; (c) coherence between CH4 and insolation; and (d) relative phasing at orbital periods. Spectral analysis followed Blackman–Tukey method, with 1/3 lags. Phases at the major orbital periods shown with 95% conﬁdence intervals.

the CH4 time scale. These age offsets lie well within the suspicious-looking distortions ofother (untuned) gas uncertainties estimated by Petit et al. (1999). signals from Vostok ice (Sections 4 and 5). The CH4 time scale produces several relatively abrupt Prior to 250,000 years ago, the ages derived from the changes in ice accumulation rate compared to the GT4 CH4 time scale diverge from those of the GT4 time scale time scale. Although these might result from ‘‘over- by much larger amounts averaging near 15,000 years tuning’’ the Vostok methane record to the July 301N (Fig. 3a). Over much ofthis interval, the GT4 time scale insolation target, none ofthese changes produces yields a climatically implausible phasing between W.F. Ruddiman, M.E. Raymo / Quaternary Science Reviews 22 (2003) 141–155 147

350

Age CH (this study) 4

300 Tuned Age (Shackleton et al., 2000)

1:1 line 250

200

150 Tuned gas ages (kyr) 100

0 0 50 100 150 200 250 300 350 (a) Glaciological (GT4) gas age (kyr)

800 600 Age GT4 (gas) July insolation (W/m July 30 N

600 500 2

400 ) Vostok methane (ppbv)

400 0 50 100 150 200 250 300 350 (b) Age (kyr)

18 Fig. 3. Comparison ofGT4, CH 4 and d Oatm time scales: (a) Diagonal (1:1) line based on Vostok GT4 gas time scale (Petit et al., 1999) is used as a 18 reference for departures of ages in other time scales. Solid line: CH4 time scale from this paper; dotted line: d Oatm time scale ofShackleton (2000); (b) Time series ofVostok CH 4 on the GT4 time scale against July insolation.

Vostok CH4 values and summer (July 301N) insolation phasing oforbital-scale ice-core signals. It also suggests (Fig. 3b): maxima in CH4 are often aligned with minima that it would be worthwhile to take a closer look at the in July insolation, a phasing opposite what would be differences between these two time scales. expected for northern hemisphere insolation control of monsoonal and boreal CH4 input. Also plotted in Fig. 3a are age-depth offsets between 18 18 the d Oatm time scale ofShackleton (2000) and the GT4 4. Assessment of the d Oatm Signal time scale. For levels prior to 250,000 years ago, the 18 18 d Oatm and CH4 time scales agree in showing ages The d Oatm time scale is based on tuning ice-core 18 considerably older than those from GT4. This conver- d Oatm to a synthetic insolation signal containing both gence supports the conclusion that the GT4 age precession and obliquity, and incorporating phase lags estimates for this older part of the Vostok record are appropriate to the slow responses ofice sheets to orbital substantially in error. forcing at these periods. In contrast, as summarized in 18 The fact that the CH4 and d Oatm time scales Section 2, the CH4 time scale is based on tuning CH4 to generally differ from each other by only a few thousand a precession-dominated 301N July insolation with no 18 years points to major progress toward the goal of lags. Although the d Oatm and CH4 time scales are obtaining a Vostok time scale sufﬁcient to analyze the based on different tuning rationales, they differ in age by 148 W.F. Ruddiman, M.E. Raymo / Quaternary Science Reviews 22 (2003) 141–155

18 less than 5000 years through most ofthe last 350,000 phase ofthe d Oatm signal through time. We compare 18 years, and by an average ofless than 500 years. the d Oatm signal (plotted on the CH4 time scale) to Despite this overall agreement, larger offsets between mid-July insolation at 301N (Fig. 4b) and to the the two time scales do occur in some intervals (Fig. 3a). SPECMAP d18O curve (Fig. 4c). 18 One clue to the origin of these offsets is that the relative The d Oatm trend shows fairly systematic phase 18 phasing ofthe CH 4 and d Oatm signals varies through offsets with respect to both insolation and SPECMAP time, with the two signals sometimes closely in phase, d18O (as also noted by Jouzel et al., 2002). During peak and sometimes well out ofphase (Fig. 4a). These shifting interglacial substages (marine isotopic stages 5.5, 9.3, 18 offsets cannot be an artifact of the time scale used: both and, to some extent, 7.3), the d Oatm signal is 18 CH4 and d Oatm are recorded in the same gas phase of approximately in phase with July insolation (Fig. 4b), the same ice core, and so similar offsets must exist for but it generally lags behind insolation during late- any time scale chosen. interglacial substages (5.1, 5.3, 7.1, and 9.1) and glacial 18 An initial way to explore the cause of these offsets is isotopic stages (4–2, 6 and 8). Conversely, d Oatm leads 18 to assume that the CH4 time scale correctly ﬁxes the age SPECMAP d O during and just after peak-interglacial ofVostok ice, and then examine resulting changes in the substages (Fig. 4c), but is roughly in phase during most

800 -0.4 )

O (‰ 0.0 Vostok CH 18 δ 600

0.4 4 (ppbv)

0.8 400 Vostok atmospheric

δ18 CH 1.2 O 4 (a) atm

-0.4 5.5 5.1 5.3 7.1 7.3 7.5 9.3 9.1 )

520 (W/m 30 July insolation N ‰ 0.0 O ( 18 δ

0.4 480

0.8 2 )

Vostok atmospheric 440

δ18 O July 30N insolation 1.2 atm (b) -0.4 5.5 5.1 9.3 5.3 2 7.1 7.3 ) 9.1

7.5 SPECMAP ‰

O ( 0.0

18 1 δ δ

0.4 18 0 ( O stack

0.8 ‰ -1 ) Vostok atmospheric

δ18O δ18 1.2 atm SPECMAP O stack -2 0 50 100 150 200 250 300 350 (c) Age (kyr)

18 Fig. 4. Evaluation ofthe Vostok d Oatm signal by time-series comparison with: (a) Vostok CH4 (both records plotted on the CH4 time scale); 18 18 (b) July 301N insolation (d Oatm plotted on the CH4 time scale and July 301N insolation from Berger and Loutre (1992); (c) d Oatm plotted on the 18 CH4 time scale and SPECMAP d O plotted on the time scale ofImbrie et al. (1984). W.F. Ruddiman, M.E. Raymo / Quaternary Science Reviews 22 (2003) 141–155 149 late-interglacial and glacial isotopic stages. These shifts inferred that the largest (estimated) Dole effect occurred 18 in d Oatm phasing appear to be linked to climate in a between peak-interglacial substages 5.5 and 5.4, ac- 18 systematic way, but why? counting for 50% or more of the d Oatm change 18 Two climatic factors control d Oatm variations over (monsoon control). The smallest Dole effect occurs in orbital time scales. The d18O value ofoxygen in the late-interglacial substages and glacial stages, where it 18 atmosphere (d Oatm) responds to the globally averaged accounts for as little as 20 to 30% of the total (ice-sheet 18 18 d O value ofmarine sea water ( d Osw), with an control). estimated lag today ofabout 1200 years caused by the We began this section by making the assumption that turnover time ofO 2 in the atmosphere (Bender et al., the ice-core methane signal has the constant phasing 18 1994). Because d Osw is a measure ofthe size ofglobal with July insolation forcing used in creating the CH4 ice sheets (and excludes the effect of ocean temperature time scale. Now we test the opposite assumption: could 18 18 on regional marine d O measurements), one ofthe two the phasing of d Oatm be constant and that ofCH 4 18 major climatic controls on d Oatm is thus global ice variable? volume. The tuning strategy used by Shackleton (2000) gave 18 The second control operates through the ‘‘Dole the d Oatm signal the same phasing as SPECMAP used effect’’ (Dole et al., 1954). Fractionation during marine for the marine d18O signal: a 5000-year lag of d18O and terrestrial photosynthesis and respiration causes behind June 21 precessional insolation and an 8000-year 18 18 modern d Oatm to be offset by +23.5% compared to lag behind obliquity. He chose smaller marine d O lags 18 d Osw. Sowers et al. (1993) found that the estimated size behind insolation than those in SPECMAP, but also 18 of this effect varied over orbital time scales by about incorporated the additional lag of d Oatm behind 18 18 18 1%, roughly the same magnitude as changes in d Osw marine d O. Because Shackleton tuned the d Oatm caused by ice sheets. Bender et al. (1994) argued that signal primarily to orbital precession, the average lag of 18 these changes in the Dole effect probably resulted from the d Oatm signal behind insolation was about 5000 varying photosynthesis and respiration in the tropical years. 18 monsoon system. Thus, changes in tropical monsoons Ifwe accept Shackleton’s choice of d Oatm phasing as 18 are the second likely climatic control on d Oatm. correct, how would it affect the timing of the CH4 18 We propose that the variable timing ofthe d Oatm signal? During and just after peak interglaciations, when 18 18 signal relative to insolation and SPECMAP d O d Oatm and CH4 are in phase, CH4 should share the (Fig. 4) is caused primarily by changes in the relative same 5000-year lag behind June 21 insolation as the 18 18 influence ofthese two controls on d Oatm through time. d Oatm signal. For the remainder ofeach interglacial- 18 During and just after peak-interglacial stages, the glacial cycle, when CH4 leads d Oatm by about 5000 18 d Oatm signal responds in phase with July insolation years, CH4 should be approximately in phase with June 18 and CH4 because the monsoon system controls d Oatm 21 insolation. in the absence oflarge ice sheets. At these times, the But this pattern ofvariable CH 4 phasing makes little 18 d Oatm signal joins CH4 in following the fast response physical sense in climatic terms. Why would a monsoon- to insolation forcing and minimal lag typical of dominated CH4 signal respond earlier when ice sheets monsoons. In addition, the rapid and early ice melting are large, but later when ice sheets are small? Ifice sheets 18 typical ofterminations brings the d Osw component of have any effect on monsoon timing, it should be to 18 the d Oatm signal into close alignment with the retard them. We conclude that the differences between 18 monsoonal Dole component. the CH4 and d Oatm time scales are more plausibly At other times, including the later parts ofinterglacial explained by variable controls on the phasing ofthe 18 18 stages and the glacial stages, the d Oatm signal responds d Oatm signal than by variable timing ofthe CH 4 signal. 18 with a phase closer to that ofmarine d Osw (ice What about the possibility that ice sheets can indeed volume), because the monsoon weakens from its peak- cause delays in monsoon timing? Ifgrowing ice sheets interglacial strength and because ice-sheet variations do retard monsoons, then the phase ofthe CH 4 signal grow large enough to become the major control ofthe should shift to younger ages during glacial isotopic 18 18 d Oatm signal. At these times, the d Oatm signal takes stages because ofthe larger lags behind insolation on the slow response to insolation changes typical ofice forcing. And if the gas-phase CH4 signal has been 18 sheets, and it lags behind the monsoon-controlled CH4 shifted in this way, the d Oatm signal must also shift to 18 signal. In summary, the phasing ofthe d Oatm signal is younger ages because it is recorded in the same bubbles. 18 monsoon-dominated when ice sheets are small, and ice- But such a shift in ice-core d Oatm creates a major dominated when ice sheets are larger. problem for its phase relationship with marine d18O 18 Other evidence supports this interpretation. The size (Fig. 4c). Any major shift of the ice-core d Oatm signal of the monsoonal Dole effect can be monitored by to younger ages will make it lag marine d18O at most 18 variations in amplitude ofthe d Oatm signal in excess of isotopic transitions. Ifthis occurs, an implausible 18 18 those in the marine d Osw signal. Sowers et al. (1993) scenario is created: the ice-core d Oatm signal now lags 150 W.F. Ruddiman, M.E. Raymo / Quaternary Science Reviews 22 (2003) 141–155 well behind its only known controlling processes: the prominent isotopic transitions occurring nearly in monsoon signal (through the Dole effect), and the phase. The largest exception is a CO2 lead of4000 to marine d18O signal. Because this scenario makes no 5000 years at the isotope stage 6/5 deglaciation sense, it seems unlikely that ice sheets can cause major (termination II). In contrast, the CO2 signal plotted on lags in the monsoon-generated CH4 signal. the GT4 time scale shows a more erratic pattern ofleads Although it is difficult (impossible?) to show conclu- and lags compared to SPECMAP d18O (Fig. 6a). sively that the monsoon-dominated CH4 signal retains This improved correlation is confirmed in a quantita- the same phasing through all the complex changes tive way by cross-spectral analysis (Figs. 5b–d and 6b-d). occurring during glacial–interglacial climate cycles, the The 23,000-year orbital period has a greatly improved evidence reviewed above suggests that it is likely that the coherency in the CH4 time scale, and the already strong CH4 signal retains a stable (or very nearly stable) phase coherence ofthe 41,000-year and 100,000-year cycles in through time. In summary, the time scales derived by the GT4 time scale both increase in the CH4 time scale. 18 tuning ice-core d Oatm and CH4 to insolation diverge Because CO2 was not used to tune the CH4 time scale, from the time scale of Petit et al. (1999) prior to 250,000 the improved coherences at all three orbital periods are years ago, but generally match each other to within a independent confirmation ofthe validity ofthe CH4 18 few thousand years back to 350,000 years ago. On time scale. In comparison, the d Oatm time scale from average, the CH4 time scale is just 450 years younger Shackleton (2000) has power at 23,000 years, but no 18 than the d Oatm time scale ofShackleton (2000). Ifthe distinct peak at that period. 18 smaller differences between the two tuned time scales are The more similar phasing ofthe CO 2 and d O signals 18 indeed caused mainly by variable climatic controls on resulting from the CH4 (and d Oatm) time scale fits into 18 d Oatm phasing, the CH4 time scale is the more accurate a long trend in the history ofVostok time-scale ofthe two. This level ofaccuracy (to within a few development. Earlier time scales based on models of thousand years) allows a close examination ofthe ice flow and accumulation (Lorius et al., 1985), and phasing ofother gases in Vostok ice. adjusted from ice age to gas age (Barnola et al., 1987), produced an irregular phasing between CO2 and marine 18 d O: large CO2 leads occurred on terminations, but 18 5. Comparison of Vostok CO2 Versus SPECMAP d O sizeable lags occurred at the marine isotopic stages 5/4 and 5e/5d transitions (times ofinferredmajor ice-sheet Ofall the signals recorded in Vostok ice, the growth). Later time scales (summarized in Section 2) concentration ofcarbon dioxide is the most important generally tended to reduce this variability in phasing, to global climate. Because the CO2 signal was not used and the CH4 time scale devised here points to a still in the tuning process, it provides an independent way to more consistent (in-phase) timing. evaluate the CH4 time scale, in this case by comparison As noted earlier, no a priori basis exists for predicting to the GT4 time scale ofPetit et al. (1999). For cross- the form of the CO2 signal. Still, it makes good sense spectral comparisons, we omit levels above 5000 years from a climate-theory viewpoint that the phasing because oflikely human influences on CH 4 (Ruddiman between CO2 and ice volume should be consistent rather and Thomson, 2001). than irregular. The relatively stable (and nearly in- 18 The concentration ofspectral power forCO 2 at phase) relationship across major d O transitions in the orbital periods is not much different for the CH4 and CH4 time scale is suggestive ofa consistent physical GT4 time scales (Figs. 5 and 6). In the CH4 time scale, relationship between CO2, a major greenhouse gas, and the peaks at 100,000 and 41,000 years lose power d18O, a primary indicator ofclimate. The increases in compared to GT4, but a new peak emerges near 23,000 coherency at all three orbital periods provide further 18 years. Also, compared to the d Oatm time scale of evidence ofa more consistent physical link between CO 2 18 Shackleton (2000), the CH4 time scale has almost and SPECMAP d O over the last 350,000 years. identical amounts ofpower at 100,000 and 41,000 years, For the CH4 time scale, maximum coherencies occur but a substantially larger peak at 23,000 years. Because at a CO2 lead of12 1 (140071400 years) relative to no a priori expectation exists about the ‘‘true’’ shape of SPECMAP d18O for the 41,000-year period and 311 the CO2 signal, and the distribution ofits power among (20007900 years) for the 23,000-year period. These 18 the orbital periods, however, these comparisons ofCO 2 phases for the d Oatm time scale agree to within 50–100 spectra alone are not a revealing test ofthe CH 4 time years. Because changes at these two periods produce scale. much ofthe distinctive character ofrapid transitions in A more diagnostic evaluation comes from comparing the SPECMAP d18O signal (Imbrie et al., 1992), the 18 the relative timing ofVostok CO 2 and the SPECMAP general visual similarity ofthe CO 2 and d O trends marine d18O signal (Imbrie et al., 1984). Plotted on the results from the similar phasing at these periods. At the 18 CH4 time scale (Fig. 5a), the CO2 signal is visually 100,000-year period, CO2 leads SPECMAP d Oby231 similar to SPECMAP d18O, with many ofthe most (660072000 years). W.F. Ruddiman, M.E. Raymo / Quaternary Science Reviews 22 (2003) 141–155 151

CO (ppmv) δ18 2 SPECMAP O 2 280 SPECMAP

1 (ppmv) 2 δ

240 18 0 ( O stack Vostok CO ‰ -1 ) 200

-2 0 50 100 150 200 250 300 350 (a) Age (kyr) kyr kyr kyr 41 23 100 4 104 6 101

CO 2

SPECMAP δ18O

BW 4 1 power 2 10 3 10 Rel.

(b) 0 100 0 100

0.8 95% CL

0.6

0.4 Coherency

0.2

Phase (deg.) -90 CO lag 2

0.00 0.05 0.10 (d) Frequency kyr-1

18 Fig. 5. Comparison ofVostok CO 2 (on the CH4 time scale) and SPECMAP d O (on the time scale ofImbrie et al., 1984). (a) Time series. (b–d) 18 Cross-spectral analysis: (b) power spectra; (c) coherency between CH4 and d O signals; and (d) relative phasing at orbital periods. Spectral analysis followed Blackman–Tukey method, with 1/3 lags. Phases at the major orbital periods shown with 95% conﬁdence intervals.

6. Discussion: Climatic role of CO2 in ice-volume changes explore this issue. On the CH4 time scale, Vostok CO2 leads SPECMAP d18O at all three orbital periods One ofthe key issues in studies oforbital-scale climate (Fig. 5). Before this can be taken as evidence that CO2 change is the phasing between CO2 and ice volume. forces ice volume, however, two problems must be Accurate definition ofthe relative phasing ofthese two addressed. signals should reveal critical cause-and-effect links in the The first problem is that, based on a wide array of climate system. The evidence above indicates that the evidence, the SPECMAP time scale is probably in error, 18 CH4 time scale for Vostok ice is sufficiently accurate to giving ages for the d O signal that average about 1500 152 W.F. Ruddiman, M.E. Raymo / Quaternary Science Reviews 22 (2003) 141–155

CO (ppmv)) 18 2 SPECMAP δ O 2 280 SPECMAP

1 (ppmv) 2 δ

240 18 0 ( O stack Vostok CO ‰ -1 ) 200

-2 0 50 100 150 200 250 300 350 (a) Age (kyr) kyr kyr kyr 41 23 100 4 104 6 101

CO2

SPECMAP δ18O

BW 4 1 power 2 10 3 10 Rel.

(b) 0 100 0 100

0.8 95% CL y 0.6

0.4 Coherenc

0.2

Phase (deg.) -90 CO lag 2

0.00 0.05 0.10 (d) Frequency kyr-1

18 Fig. 6. Comparison ofVostok CO 2 (on the GT4 time scale) and SPECMAP d O (on the time scale ofImbrie et al., 1984): (a) Time series; (b and d) 18 Cross-spectral analysis: (b) power spectra; (c) coherency between CH4 and d O signals; and (d) relative phasing at orbital periods. Spectral analysis followed Blackman–Tukey method, with 1/3 lag. Phases at the major orbital periods shown with 95% confidence intervals. years too young. Well-dated coral reefs provide a direct A second line ofevidence emerged froman attempt index ofpast ice-volume changes at some key isotopic to tune the d18O signal using a different rationale. transitions. By dating reefs formed at the end of Pisias et al. (1990) allowed the lags ofthe 41,000-year termination II and during other high stands ofsea level, and 23,000-year d18O responses behind the insolation Edwards et al. (1987) and Bard et al. (1990) found forcing signals to vary through time in order to evidence for faster deglacial sea-level rises (ice-volume optimize matches ofthe resulting d18O signal to decreases) than implied by the SPECMAP d18O curve. those observed in marine records. This approach Both concluded that coral-reefdata indicate deglacia- contrasts with the constant phase lags assumed by tion some 3000 years earlier than the timing ofthe SPECMAP in the tuning process (Imbrie et al., 1984). SPECMAP d18O signal across termination II. Pisias et al. (1990) found that on average the d18O signal W.F. Ruddiman, M.E. Raymo / Quaternary Science Reviews 22 (2003) 141–155 153 was shifted earlier by 1500 years compared to the necessarily indicate the phasing ofCO 2 with respect to SPECMAP time scale. The largest offset from SPEC- true ice volume. MAP, 3500 years on termination II, agrees with the Shackleton (2000) challenged the assumption that the coral-reefdates. 100,000-year component of d18O signals is an accurate A third line ofevidence comes froman indirect record ofice volume. He found that halfofthe 100,000- 18 comparison ofthe CH 4 time scale to the SPECMAP year d O signal in Pacific Ocean benthic foraminifera 18 18 d O time scale using the d Oatm time scale of could be explained by changes in Pacific deep-water Shackleton (2000) as an intermediary. Shackleton temperature, and the other halfby ice volume. Within 18 (2000) reported that his d Oatm time scale was on the sizeable errors ofhis multi-step analysis, the observed average 2000 years older (earlier) than that ofSPEC- phase ofthe 100,000-year d18O signal in this core turned MAP, whereas we find that the CH4 time scale averages out to be partitioned into an early phased temperature 18 450 years younger than his d Oatm time scale. Combin- component and a late ice-volume component. ing these results, the CH4 time scale presented here is The same problem complicates our attempt to 18 consistent with a 1550-year shift of the SPECMAP d O compare the phasing ofCO 2 against ice volume based time scale toward older (earlier) ages, as suggested by on the SPECMAP d18O record. The average glacial- Pisias et al. (1990). interglacial d18O amplitude ofthe cores used in the An array ofevidence thus suggests that the SPEC- SPECMAP stack is 1.7% (Imbrie et al., 1984). If0.9– MAP d18O signal lags true ice volume by as much as 1.3% ofthis total is caused by changes in ice volume 3000 years or more during rapid deglaciations and by an (Schrag et al., 1996; Shackleton, 1967), a residual signal average ofabout 1500 years throughout the record. ofbetween 0.4 % and 0.8% remains, equivalent to an With the SPECMAP time scale adjusted by 1500 years SST change of1.7–3.3 1C. Independent estimates ofSST toward older ages, the 1400-year CO2 lead relative to changes in the tropical and Southern Ocean cores from SPECMAP d18O (and inferred ice volume) at the which the SPECMAP d18O stack was compiled are 41,000-year period disappears entirely, as does much consistent with a temperature overprint ofthis size. ofthe 2000-year lead at the 23,000-year period. CO 2 Because an SST overprint of0.4–0.7 % is larger than retains a lead of5000 years at the 100,000-year period of either the 23,000-year or 41,000-year components ofthe orbital eccentricity. SPECMAP d18O signal, and amounts to halfor more of 18 18 IfCO 2 is nearly in phase with d O (ice volume) at a the filtered 100,000-year d O component (Imbrie et al., given orbital period, it must be acting mainly as a 1989), it could have a major effect on the phase of any or positive (amplifying) feedback to the ice sheets, rather all ofthese d18O signals. SPECMAP concluded that SST than as an independent source offorcing. This conclu- leads d18O at all three orbital periods in the tropical and sion derives from the long time constant of ice-sheet Southern Ocean region ofthe SPECMAP cores (Imbrie responses (Weertman; 1964; Imbrie et al., 1984). Ice et al., 1992, 1993). As a result, any SST overprint in this volume should lag thousands ofyears behind its primary region would probably be phased ahead ofthe d18O source(s) of forcing. Ice sheets cannot be strongly forced signals observed. Ifthis is the case, the phase ofthe true 18 by a CO2 signal that has nearly the same phasing, except ice-volume signal would have to lag behind d Oto in the sense of receiving a self-amplifying positive balance the early SST overprint. This means that CO2 is feedback. The close phase relationship between CO2 likely to lead ice volume by at least as much as it leads and d18O at the 41,000-year obliquity cycle is consistent d18O, and possibly by more. with this kind offeedbackrelationship. In contrast, the Without a much more detailed analysis, the over- 18 18 CO2 leads versus d O at the 100,000-year cycles (and to printing effect of SST on d O at each orbital period a lesser extent the 23,000-year cycle) still permit an cannot be deciphered. We can rule out the possibility 18 active forcing role for CO2,ifd O is an accurate index that CO2 lags behind ice volume at any ofthe orbital ofice volume. periods. But we cannot distinguish between the possi- 18 Unfortunately, marine d O signals are not exact bility that CO2 leads ice volume, in which case it is an proxies for ice volume. The major complication is that active part ofthe forcing ofthe ice sheets, or is in phase marine d18O records also contain local temperature with ice volume, in which case it plays the role ofan signals. Estimates ofthe fractionofmarine d18O signals amplifying positive feedback to ice-volume changes. caused by temperature changes range from 1/3 (Shack- leton, 1967) to 1/2 (Schrag et al., 1996), and the effect of the temperature overprints on d18O phasing is unclear Acknowledgements (Mix et al., 2001). A second complication is that changes in d18O composition ofice sheets through time may M.E. Raymo thanks NSF for the research support produce offsets between marine d18O and ice volume provided by the MGG and ODP panels and specifically (Mix and Ruddiman, 1984). As a result, the phasing of a generous supplement to grant OCE-9631759. 18 CO2 with respect to the marine d O signal does not W. Ruddiman thanks his slowly declining TIAA/CREF 154 W.F. Ruddiman, M.E. Raymo / Quaternary Science Reviews 22 (2003) 141–155 investments for financial support. We also thank Alan Yiou, F., Yiou, P., 1993. Extending the Vostok ice-core record of Mix and an anonymous reviewer for careful and paleoclimate to the penultimate glacial period. Nature 364, 407–412. comprehensive reviews ofthe initial manuscript. Jouzel, J., Hoffman, G., Parrenin, F., Waelbroeck, C., 2002. Atmospheric 18O and sea level. Quaternary Science Reviews 21, 307–314. References Kutzbach, J.E., 1981. Monsoon climate ofthe early Holocene: climate experiment with earth’s orbital parameters for 9000 years ago. Bard, E., Hamelin, B., Fairbanks, R.F., 1990. U–Th ages obtained by Science 214, 59–61. mass spectrometry in corals from Barbados: sea level during the Kutzbach, J.E., Street-Perrott, F.A., 1985. Milankovitch forcing of past 130,000 years. Nature 346, 456–458. fluctuations in the level oftropical lakes from18 to 0 kyr BP. Barnola, J.M., Raynaud, D., Korotkevitch, Y.S., Lorius, C., 1987. Nature 317, 130–134. Vostok ice cores provides 160,000-year record ofatmospheric CO 2. Lorius, C., Jouzel, J., Ritz, C., Merlivat, L., Barkov, N.I., Nature 329, 408–414. Korotkevtich, Y.S., Kotlyakov, V.M., 1985. A 150,000-year Bender, M., Sowers, T., Labeyrie, L., 1994. The Dole effect and its climatic record from Antarctic ice. Nature 316, 591–596. variations during the last 130,000 years as measured in the Vostok Meese, D.A., Gow, A.J., Alley, R.B., Zielinski, G.A., Grootes, P.M., ice core. Global Biogeochemical Cycles 8, 363–376. Ram, K.C., Taylor, K.C., Mayewski, P.A., Bolzan, J.F., 1997. The Berger, A., Loutre, M.F., 1992. Astronomical solutions for paleocli- Greenland ice sheet project 2 depth-age scale: methods and results. mate studies over the last 3 million years. Earth and Planetary Journal Geophysical Research 102, 26411–26423. Science letters 111, 369–392. McIntyre, A., Ruddiman, W.F., Karlin, K., Mix, A.C., 1989. Surface Blunier, T., Chappellaz, J., Schwander, J., Stauffer, J., Raynaud, D., water response ofthe equatorial Atlantic Ocean to orbital forcing. 1995. Variations in atmospheric methane concentration during the Paleoceanography 4, 19–55. Holocene epoch. Nature 374, 46–49. Mix, A.C., Ruddiman, W.F., 1984. Oxygen-isotope analyses and Brook, E.J., Sowers, T., Orchado, J., 1996. Rapid variations in Pleistocene ice volume. Quaternary Research 21, 1–20. atmospheric methane concentration during the past 110,000 years. Mix, A., Bard, E., Schneider, R., 2001. Environmental processes ofthe Science 273, 1087–1091. ice age: land, oceans, glaciers (EPILOG). Quaternary Science Brook, E.J., Harder, S., Severinghaus, J., Steig, E., Sucker, C.M., Reviews 20, 627–657. 2000. On the origin and timing ofrapid changes in atmospheric Petit, J.R., Jouzel, J., Raynaud, D., Barkov, N.I., Barnola, J.-M., methane during the last glacial period. Global Biogeochemical Basile, I., Bender, M., Chappellaz, J., Davis, M., Delaygue, G., Cycles 14, 559–572. Delmotte, M., Kotlyakov, V.M., Lipenkov, V., Lorius, C., Pepin, Chappellaz, J., Barnola, J.-M., Raynaud, D., Korotkevitch, Y.S., L., Ritz, C., Saltzman, E., Stievenard, M., 1999. Climate and Lorius, C., 1990. Atmospheric CH record over the last climatic 4 atmospheric history ofthe last 420,000 years fromthe Vostok ice cycle revealed by the Vostok ice core. Nature 345, 127–131. core, Antarctica. Nature 399, 429–436. Chappellaz, J., Blunier, T., Kints, S., Dallenbach, A., Barnola, J.-M., Pichon, J.-J., Labeyrie, L.D., Bareille, G., Labracherie, M., Duprat, J., Schwander, J., Raynaud, D., Stauffer, B., 1997. Change in the Jouzel, J., 1992. Surface water temperature changes in the high atmospheric CH gradient between Greenland and Antarctica 4 latitudes ofthe southern hemisphere over the last glacial– during the Holocene. Journal ofGeophysical Research 102 (D), interglacial cycle. Paleoceanography 7, 289–318. 15987–15997. Pisias, N.G., Mix, A.C., Zahn, R., 1990. Nonlinear response in the deMenocal, P.B., Rind, D., 1996. Sensitivity ofsubtropical African global climate system: evidence from benthic oxygen isotopic and Asian climate to prescribed boundary condition changes: record in core RC13-110. Paleoceanography 5, 147–160. model implications for the Plio–Pleistocene evolution of low- Pokras, E.M., Mix, A.C., 1987. Earth’s precession cycle and Quaternary latitude climate. In: Johnson, T.C., Okada, E.O (Eds.), The Limnology, Climatology, and Paleoclimatology ofthe East African climatic change in tropical Africa. Nature 326, 486–487. Lakes. Gordon and Breach, Amsterdam, pp. 57–77. Prell, W.L., 1984. Monsoonal climate ofthe Arabian Sea during the Dole, M., Lane, G.A., Rudd, D.P., Zuakelies, D.A., 1954. Isotopic late Quaternary: a response to changing solar radiation. In: Berger, composition ofatmospheric oxygen and nitrogen. Geochimica A.L., et al. (Ed.), Milankovitch and Climate. Reidel, Dordrecht, Cosmochimica Acta 6, 65–78. pp. 349–366. Edwards, R.L., Chen, J.H., Ku, T.-L., Wasserburg, G.J., 1987. Precise Prell, W.L., van Campo, E., 1986. Coherent response ofArabian Sea timing ofthe last interglacial period frommass-spectrometric upwelling and pollen transport to late quaternary monsoonal determination of 230Th in corals. Science 236, 1547–1553. winds. Nature 323, 526–528. Gasse, F., 2000. Hydrological changes in the African tropics since the Prell, W.L., Kutzbach, J.E., 1987. Monsoon variability over the last last glacial maximum. Quaternary Science Reviews 19, 189–211. 150,000 years. Journal Geophysical Research 92, 8411–8425. Imbrie, J., et al., 1984. The orbital theory ofPleistocene climate: Raymo, M.E., Horowitz, M., 1996. Organic carbon paleo-pCO2 and support from a revised chronology of the marine d18O record. In: marine-ice core correlations. Geophysical Research Letters 23, Berger, A.L., et al. (Ed.), Milankovitch and Climate, Part I. Riedel, 367–370. Dordrecht, pp. 269–305. Raynaud, D., Chappellaz, J., Barnola, J.M., Korotkevitch, Y.S., Imbrie, J., McIntyre, A., Mix, A., 1989. Oceanic response to orbital Lorius, C., 1988. Climatic and CH4 cycle implications ofglacial- forcing in the late Quaternary: observational and experimental interglacial CH4 change in the Vostok ice core. Nature 333, strategies. In: Berger, A., et al. (Ed.), Climate and Geo-Sciences. 655–657. Kluwer Academic Publications, Dordrecht, pp. 121–164. Rossignol-Strick, M., Nesteroff, W., Olive, P., Vergnaud-Grazzini, C., Imbrie, J., et al., 1992. On the structure and origin ofmajor glaciation 1982. After the deluge: Mediterranean stagnation and sapropel cycles. I. Linear responses to Milankovitch forcing. Paleoceano- formation. Nature 295, 105–110. graphy 7, 701–738. Ruddiman, W.F., Thomson, J.S., 2001. The case for human causes of Imbrie, J., et al., 1993. On the structure and origin ofmajor glaciation increased atmospheric CH4 over the last 5000 years. Quaternary cycles. II. The 100,000-year cycle. Paleoceanography 8, 699–735. Science Reviews 20, 1769–1777. Jouzel, J., Barkov, N.I., Barnola, J.M., Bender, M., Chappellaz, J., Schrag, D., Hampt, G., Murray, D.W., 1996. Pore fluid constraints on Genthon, C., Kotlyakov, V.M., Lipenkov, V., Lorius, C., Petit, J.R., the temperature and oxygen isotopic composition ofthe glacial Raynaud, D., Raisbeck, G., Ritz, C., Sowers, T., Stievenard, M., ocean. Science 272, 1930–1932. W.F. Ruddiman, M.E. Raymo / Quaternary Science Reviews 22 (2003) 141–155 155

Shackleton, N.J., 1967. Oxygen isotope analyses and Pleistocene Thunell, R.C., Williams, D.F., 1989. Glacial-Holocene salinity changes temperatures re-assessed. Nature 215, 15–17. in the Mediterranean Sea: hydrographic and depositional effects. Shackleton, N.J., 2000. The 100,000-year ice-age cycle identiﬁed and Nature 338, 493–496. found to lag temperature, carbon dioxide, and orbital eccentricity. Waelbroeck,C.,Jouzel,J.,Labeyrie,L.,Lorius,C.,Labracherie,M., Science 289, 1897–1902. Stievenard, M., Barkov, N., 1995. A comparison ofthe Vostok ice de- Sowers, T., Bender, M., Labeyrie, L., Martinson, D., Jouzel, J., uterium record and series from Southern Ocean core MD 88-70 over Raynaud, D., Pichon, J.J., Korotkevitch, Y.S., 1993. A 135,000- the last two glacial-interglacial cycles. Climate Dynamics 12, 113–123. year Vostok-SPECMAP common temporal scale. Paleoceanogra- Wang, L., Sarnthein, M., Erlenkeuser, H., Heilig, S., Lienast, M., phy 8, 737–766. Pﬂaumann, U., Pelejero, C., Grootes, P.M., 1999. East Asian Street-Perrott, F.A., Harrison, S.P., 1984. Temporal variations in lake monsoon climate history during the late Quaternary: high levels since 30,000 years BP-an index ofthe global hydrological resolution sediment records from the south China sea. Marine cycle. In: Hansen, J.E., Takahashi, T. (Eds.), Climate Processes and Geology 156, 245–284. Climate Sensitivity, Geophysical Monograph Series, Vol. 29. Weertman, J., 1964. Rate ofgrowth or shrinkage ofnon-equilibrium AGU, Washington, DC, pp. 118–129. ice sheets. Journal Glaciology 5, 145–158.