# Climate and atmospheric history of the past 420,000 years from the Vostok ice core, Antarctica

Jean-Robert Petit, Jean Jouzel, Dominique Raynaud, Nartsiss Irinarkhovich Barkov, Jean-Marc Barnola, Isabelle Basile-Doelsch, Michael Bender, Jérôme Chappellaz, Mike Davis, Gilles Delaygue, et al.

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Jean-Robert Petit, Jean Jouzel, Dominique Raynaud, Nartsiss Irinarkhovich Barkov, Jean-Marc Barnola, et al. Climate and atmospheric history of the past 420,000 years from the Vostok ice core, Antarctica. Nature, 1999, 399 (6735), pp.429-436. 10.1038/20859. hal-00756651

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Distributed under a Creative Commons Attribution 4.0 International License
**Title**

Climate and atmospheric history of the past 420,000 years from the Vostok ice core, Antarctica

**Permalink**

<https://escholarship.org/uc/item/7rx4413n>

**Journal**

Nature, 399(6735)

**ISSN**

0028-0836

**Authors**

Petit, JR
Jouzel, J
Raynaud, D
et al.

**Publication Date**

1999-06-01

**DOI**

10.1038/20859

**Supplemental Material**

<https://escholarship.org/uc/item/7rx4413n#supplemental>

**Copyright Information**

This work is made available under the terms of a Creative Commons Attribution License, available at <https://creativecommons.org/licenses/by/4.0/>

Peer reviewed
# Climate and atmospheric history of the past 420,000 years from the Vostok ice core, Antarctica

J. R. Petit*, J. Jouzel†, D. Raynaud*, N. I. Barkov‡, J.-M. Barnola*, I. Basile*, M. Bender§, J. Chappellaz*, M. Davis||, G. Delaygue†, M. Delmotte*, V. M. Kotlyakov¶, M. Legrand*, V. Y. Lipenkov‡, C. Lorius*, L. Pépin*, C. Ritz*, E. Saltzman|| & M. Stievenard†

**The recent completion of drilling at Vostok station in East Antarctica has allowed the extension of the ice record of atmospheric composition and climate to the past four glacial-interglacial cycles. The succession of changes through each climate cycle and termination was similar, and atmospheric and climate properties oscillated between stable bounds. Interglacial periods differed in temporal evolution and duration. Atmospheric concentrations of carbon dioxide and methane correlate well with Antarctic air-temperature throughout the record. Present-day atmospheric burdens of these two important greenhouse gases seem to have been unprecedented during the past 420,000 years.**

The late Quaternary period (the past one million years) is punctuated by a series of large glacial-interglacial changes with cycles that last about 100,000 years (ref. 1). Glacial-interglacial climate changes are documented by complementary climate records¹,² largely derived from deep sea sediments, continental deposits of flora, fauna and loess, and ice cores. These studies have documented the wide range of climate variability on Earth. They have shown that much of the variability occurs with periodicities corresponding to that of the precession, obliquity and eccentricity of the Earth's orbit¹,³. But understanding how the climate system responds to this initial orbital forcing is still an important issue in palaeoclimatology, in particular for the generally strong ~100,000-year (100-kyr) cycle.

Ice cores give access to palaeoclimate series that includes local temperature and precipitation rate, moisture source conditions, wind strength and aerosol fluxes of marine, volcanic, terrestrial, cosmogenic and anthropogenic origin. They are also unique with their entrapped air inclusions in providing direct records of past changes in atmospheric trace-gas composition. The ice-drilling project undertaken in the framework of a long-term collaboration between Russia, the United States and France at the Russian Vostok station in East Antarctica (78°S, 106°E, elevation 3,488 m, mean temperature -55°C) has already provided a wealth of such information for the past two glacial-interglacial cycles⁴⁻¹³. Glacial periods in Antarctica are characterized by much colder temperatures, reduced precipitation and more vigorous large-scale atmospheric circulation. There is a close correlation between Antarctic temperature and atmospheric concentrations of CO₂ and CH₄ (refs 5, 9). This discovery suggests that greenhouse gases are important as amplifiers of the initial orbital forcing and may have significantly contributed to the glacial-interglacial changes¹⁴⁻¹⁶. The Vostok ice cores were also used to infer an empirical estimate of the sensitivity of global climate to future anthropogenic increases of greenhouse-gas concentrations¹⁵.

The recent completion of the ice-core drilling at Vostok allows us to considerably extend the ice-core record of climate properties at this site. In January 1998, the Vostok project yielded the deepest ice

core ever recovered, reaching a depth of 3,623 m (ref. 17). Drilling then stopped ~120 m above the surface of the Vostok lake, a deep subglacial lake which extends below the ice sheet over a large area¹⁸, in order to avoid any risk that drilling fluid would contaminate the lake water. Preliminary data¹⁷ indicated that the Vostok ice-core record extended through four climate cycles, with ice slightly older than 400 kyr at a depth of 3,310 m, thus spanning a period comparable to that covered by numerous oceanic¹ and continental² records.

Here we present a series of detailed Vostok records covering this ~400-kyr period. We show that the main features of the more recent Vostok climate cycle resemble those observed in earlier cycles. In particular, we confirm the strong correlation between atmospheric greenhouse-gas concentrations and Antarctic temperature, as well as the strong imprint of obliquity and precession in most of the climate time series. Our records reveal both similarities and differences between the successive interglacial periods. They suggest the lead of Antarctic air temperature, and of atmospheric greenhouse-gas concentrations, with respect to global ice volume and Greenland air-temperature changes during glacial terminations.

## The ice record

The data are shown in Figs 1, 2 and 3 (see Supplementary Information for the numerical data). They include the deuterium content of the ice (δDice, a proxy of local temperature change), the dust content (desert aerosols), the concentration of sodium (marine aerosol), and from the entrapped air the greenhouse gases CO₂ and CH₄, and the δ¹⁸O of O₂ (hereafter δ¹⁸Oatm) which reflects changes in global ice volume and in the hydrological cycle¹⁹. (δD and δ¹⁸O are defined in the legends to Figs 1 and 2, respectively.) All these measurements have been performed using methods previously described except for slight modifications (see figure legends).

The detailed record of δDice (Fig. 1) confirms the main features of the third and fourth climate cycles previously illustrated by the coarse-resolution record¹⁷. However, a sudden decrease from interglacial-like to glacial-like values, rapidly followed by an abrupt return to interglacial-like values, occurs between 3,320 and 3,330 m.

<footnote>* Laboratoire de Glaciologie et Géophysique de l'Environnement, CNRS, BP96, 38402, Saint Martin d'Hères Cedex, France</footnote>

<footnote>† Laboratoire des Sciences du Climat et de l'Environnement (UMR CEA/CNRS 1572), L'Orme des Merisiers, Bât. 709, CEA Saclay, 91191 Gif-sur-Yvette Cedex, France</footnote>

<footnote>‡ Arctic and Antarctic Research Institute, Beringa Street 38, 199397, St Petersburg, Russia</footnote>

<footnote>§ Department of Geosciences, Princeton University, Princeton, New Jersey 08544-1003, USA</footnote>

<footnote>|| Rosenstiel School of Marine and Atmospheric Science, University of Miami, 4600 Rickenbacker Causeway, Miami, Florida 33149, USA</footnote>

<footnote>¶ Institute of Geography, Staromonety, per 29, 109017, Moscow, Russia</footnote>
In addition, a transition from low to high CO₂ and CH₄ values (not shown) occurs at exactly the same depth. In undisturbed ice, the transition in atmospheric composition would be found a few metres lower (due to the difference between the age of the ice and the age of the gas²⁰). Also, three volcanic ash layers, just a few centimetres apart but inclined in opposite directions, have been observed—10 m

above this δD excursion (3,311 m). Similar inclined layers were observed in the deepest part of the GRIP and GISP2 ice cores from central Greenland, where they are believed to be associated with ice flow disturbances. Vostok climate records are thus probably disturbed below these ash layers, whereas none of the six records show any indication of disturbances above this level. We therefore limit

<figure><img src="image_2.png" /></figure>

**Figure 1** The deuterium record. Deuterium content as a function of depth, expressed as δD (in ‰ with respect to Standard Mean Ocean Water, SMOW). This record combines data available down to 2,755 m (ref. 13) and new measurements performed on core 5G (continuous 1-m ice increments) from 2,755 m to 3,350 m.

Measurement accuracy (1σ) is better than 1‰. Inset, the detailed deuterium profile for the lowest part of the record showing a δD excursion between 3,320 and 3,330 m. $\delta D_{\text{ice}}(\text{in ‰}) = [(D/H)_{\text{sample}}/(D/H)_{\text{SMOW}} - 1] \times 1,000$.

<figure><img src="image_3.png" /></figure>

**Figure 2** Vostok time series and ice volume. Time series (GT4 timescale for ice on the lower axis, with indication of corresponding depths on the top axis and indication of the two fixed points at 110 and 390 kyr) of: **a**, deuterium profile (from Fig. 1); **b**, $\delta^{18}\text{O}_{\text{atm}}$ profile obtained combining published data¹¹,¹³,³⁰ and 81 new measurements performed below 2,760 m. The age of the gas is calculated as described in ref. 20; **c**, seawater $\delta^{18}\text{O}$ (ice volume proxy) and marine isotope stages adapted from Bassinot et al.²⁶; **d**, sodium profile obtained by combination

of published and new measurements (performed both at LGGE and RSMAS) with a mean sampling interval of 3–4 m (ng g⁻¹ or p.p.b); and **e**, dust profile (volume of particles measured using a Coulter counter) combining published data¹⁰,¹³ and extended below 2,760 m, every 4 m on the average (concentrations are expressed in µg g⁻¹ or p.p.m. assuming that Antarctic dust has a density of 2,500 kg m⁻³). $\delta^{18}\text{O}_{\text{atm}}(\text{in ‰}) = [(^{18}\text{O}/^{16}\text{O})_{\text{sample}}/(^{18}\text{O}/^{16}\text{O})_{\text{standard}} - 1] \times 1,000$; standard is modern air composition.
the discussion of our new data sets to the upper 3,310 m of the ice core, that is, down to the interglacial corresponding to marine stage 11.3.

Lorius *et al.*⁴ established a glaciological timescale for the first climate cycle of Vostok by combining an ice-flow model and an ice-accumulation model. This model was extended and modified in several studies¹²,¹³. The glaciological timescale provides a chronology based on physics, which makes no assumption about climate forcings or climate correlation except for one or two adopted control ages. Here, we further extend the Extended Glaciological Timescale (EGT) of Jouzel *et al.*¹² to derive GT4, which we adopt as our primary chronology (see Box 1). GT4 provides an age of 423 kyr at a depth of 3,310 m.

## Climate and atmospheric trends

**Temperature.** As a result of fractionation processes, the isotopic content of snow in East Antarctica (δD or δ¹⁸O) is linearly related to the temperature above the inversion level, Tᵢ, where precipitation forms, and also to the surface temperature of the precipitation site, Tₛ (with ΔT₁ = 0.67ΔTₛ, see ref. 6). We calculate temperature changes from the present temperature at the atmospheric level as ΔTᵢ = (ΔδDᵢice - 8Δδ¹⁸Oₛw)/9, where Δδ¹⁸Oₛw is the globally averaged change from today's value of seawater δ¹⁸O, and 9‰ per °C is the spatial isotope/temperature gradient derived from deuterium data in this sector of East Antarctica²¹. We applied the above relationship to calculate ΔTₛ. This approach underestimates ΔTₛ by a factor of ~2 in Greenland²² and, possibly, by up to 50% in Antarctica²³. However, recent model results suggest that any underestimation of temperature changes from this equation is small for Antarctica²⁴,²⁵.

To calculate ΔT₁ from δD, we need to adopt a curve for the change in the isotopic composition of sea water versus time and correlate it with Vostok. We use the stacked δ¹⁸Oₛw record of Bassinot *et al.*²⁶, scaled with respect to the V19-30 marine sediment record over their common part that covers the past 340 kyr (ref. 27) (Fig. 2). To avoid distortions in the calculation of ΔT₁ linked with dating uncertainties, we correlate the records by performing a peak to peak adjustment between the ice and ocean isotopic records. The δ¹⁸Oₛw correction corresponds to a maximum ΔT₁ correction of ~1°C and associated uncertainties are therefore small. We do not attempt to correct ΔT₁ either for the change of the altitude of the ice sheet or for the origin of the ice upstream of Vostok¹³; these terms are very poorly known and, in any case, are also small (<1°C).

The overall amplitude of the glacial-interglacial temperature change is ~8°C for ΔT₁ (inversion level) and ~12°C for ΔTₛ, the temperature at the surface (Fig. 3). Broad features of this record are thought to be of large geographical significance (Antarctica and part of the Southern Hemisphere), at least qualitatively. When examined in detail, however, the Vostok record may differ from coastal²⁸ sites in East Antarctica and perhaps from West Antarctica as well.

Jouzel *et al.*¹³ noted that temperature variations estimated from deuterium were similar for the last two glacial periods. The third and fourth climate cycles are of shorter duration than the first two cycles in the Vostok record. The same is true in the deep-sea record, where the third and fourth cycles span four precessional cycles rather than five as for the last two cycles (Fig. 3). Despite this difference, one observes, for all four climate cycles, the same 'sawtooth' sequence of a warm interglacial (stages 11.3, 9.3, 7.5 and 5.5), followed by increasingly colder interstadial events, and ending with a rapid return towards the following interglacial. The

<figure><img src="image_4.png" /></figure>

**Figure 3** Vostok time series and insolation. Series with respect to time (GT4 timescale for ice on the lower axis, with indication of corresponding depths on the top axis) of: a, CO₂; b, isotopic temperature of the atmosphere (see text); c, CH₄; d, δ¹⁸O_atm; and e, mid-June insolation at 65°N (in W m⁻²) (ref. 3). CO₂ and CH₄ measurements have been performed using the methods and analytical procedures previously described⁵,⁹. However, the CO₂ measuring system has been slightly modified in order to increase the sensitivity of the CO₂ detection. The

thermal conductivity chromatographic detector has been replaced by a flame ionization detector which measures CO₂ after its transformation into CH₄. The mean resolution of the CO₂ (CH₄) profile is about 1,500 (950) years. It goes up to about 6,000 years for CO₂ in the fractured zones and in the bottom part of the record, whereas the CH₄ time resolution ranges between a few tens of years to 4,500 years. The overall accuracy for CH₄ and CO₂ measurements are ±20 p.p.b.v. and 2-3 p.p.m.v., respectively. No gravitational correction has been applied.
coolest part of each glacial period occurs just before the glacial termination, except for the third cycle. This may reflect the fact that the June 65°N insolation minimum preceding this transition (255 kyr ago) has higher insolation than the previous one (280 kyr ago), unlike the three other glacial periods. Nonetheless, minimum

Box 1 The Vostok glaciological timescale

We use three basic assumptions¹² to derive our glaciological timescale (GT4); (1) the accumulation rate has in the past varied in proportion to the derivative of the water vapour saturation pressure with respect to temperature at the level where precipitation forms (see section on the isotope temperature record), (2) at any given time the accumulation between Vostok and Dome B (upstream of Vostok) varies linearly with distance along the line connecting those two sites, and (3) the Vostok ice at 1,534 m corresponds to marine stage 5.4 (110 kyr) and ice at 3,254 m corresponds to stage 11.2.4 (390 kyr).

Calculation of the strain-induced thinning of annual layers is now performed accounting for the existence of the subglacial Vostok lake. Indeed, running the ice-flow model⁴⁸ with no melting and no basal sliding as done for EGT¹² leads to an age >1,000 kyr for the deepest level we consider here (3,310 m), which is much too old. Instead, we now allow for moderate melting and sliding. These processes diminish thinning for the lower part of the core and provide younger chronologies. We ran this age model⁴⁸ over a large range of values of the model parameters (present-day accumulation at Vostok, *A*, melting rate, *M*, and fraction of horizontal velocity due to base sliding, *S*) with this aim of matching the assumed ages at 1,534 and 3,254 m. This goal was first achieved (ages of 110 and 392 kyr) with $A = 1.96 \text{ g cm}^{-2} \text{ yr}^{-1}$, and *M* and *S* equal respectively to 0.4 mm yr⁻¹ and 0.7 for the region 60 km around Vostok where the base is supposed to reach the melting point (we set *M* = 0 and *S* = 0 elsewhere). These values are in good agreement with observations for *A* (2.00 ± 0.04 g cm⁻² yr⁻¹ over the past 200 yr) and correspond to a reasonable set of parameters for *M* and *S*. We adopt this glaciological timescale (GT4), which gives an age of 423 kyr at 3,310 m, without further tuning (Fig. 2). GT4 never differs by more than 2 kyr from EGT over the last climate cycle and, in qualitative agreement with recent results⁴⁹, makes termination I slightly older (by ~700 yr). We note that it provides a reasonable age for stage 7.5 (238 kyr) whereas Jouzel *et al.*¹³ had to modify EGT for the second climate cycle by increasing the accumulation by 12% for ages older than 110 kyr. GT4 never differs by more than 4 kyr from the orbitally tuned timescale of Waelbroeck *et al.*⁵⁰ (defined back to 225 kyr), which is within the estimated uncertainty of this latter timescale. Overall, we have good arguments¹¹,⁵⁰⁻⁵² to claim that the accuracy of GT4 should be better than ±5 kyr for the past 110 kyr.

The strong relationship between δ¹⁸O_atm and mid-June 65°N insolation changes (see text and Fig. 3) enables us to further evaluate the overall quality of GT4. We can use each well-marked transition from high to low δ¹⁸O_atm to define a 'control point' giving an orbitally tuned age. The mid-point of the last δ¹⁸O_atm transition (~10 kyr ago) has nearly the same age as the insolation maximum (11 kyr). We assume that this correspondence also holds for earlier insolation maxima. The resulting control points (Fig. 3 and Table 1) are easy to define for the period over which the precessional cycle is well imprinted in 65°N insolation (approximately between 60 and 340 kyr) but not during stages 2 and 10 where insolation changes are small. The agreement between the δ¹⁸O_atm control points and GT4 is remarkably good given the simple assumptions of both approaches. This conclusion stands despite the fact that we do not understand controls on δ¹⁸O_atm sufficiently well enough to know about the stability of its phase with respect to insolation. We assume that the change in phase does not exceed ±6 kyr (1/4 of a precessional period).

We conclude that accuracy of GT4 is always better than ±15 kyr, better than ±10 kyr for most of the record, and better than ±5 kyr for the last 110 kyr. This timescale is quite adequate for the discussions here which focus on the climatic information contained in the Vostok records themselves.

temperatures are remarkably similar, within 1°C, for the four climate cycles. The new data confirm that the warmest temperature at stage 7.5 was slightly warmer than the Holocene¹³, and show that stage 9.3 (where the highest deuterium value, -414.8‰, is found) was at least as warm as stage 5.5. That part of stage 11.3, which is present in Vostok, does not correspond to a particularly warm climate as suggested for this period by deep-sea sediment records²⁹. As noted above, however, the Vostok records are probably disturbed below 3,310 m, and we may not have sampled the warmest ice of this interglacial. In general, climate cycles are more uniform at Vostok than in deep-sea core records¹. The climate record makes it unlikely that the West Antarctic ice sheet collapsed during the past 420 kyr (or at least shows a marked insensitivity of the central part of East Antarctica and its climate to such a disintegration).

The power spectrum of ΔT₁ (Fig. 4) shows a large concentration of variance (37%) in the 100-kyr band along with a significant concentration (23%) in the obliquity band (peak at 41 kyr). This strong obliquity component is roughly in phase with the annual insolation at the Vostok site⁴,⁶,¹⁵. The variability of annual insolation at 78°S is relatively large, 7% (ref. 3). This supports the notion that annual insolation changes in high southern latitudes influence Vostok temperature¹⁵. These changes may, in particular, contribute to the initiation of Antarctic warming during major terminations, which (as we show below) herald the start of deglaciation.

There is little variance (11%) in ΔT₁ around precessional periodicities (23 and 19 kyr). In this band, the position of the spectral peaks is affected by uncertainties in the timescale. To illustrate this point, we carried out, as a sensitivity test, a spectral analysis using the control points provided by the δ¹⁸O_atm record (see Table 1). The position and strength of the 100- and 40-kyr-spectral peaks are unaffected, whereas the power spectrum is significantly modified for periodicities lower than 30 kyr.

**Insolation.** δ¹⁸O_atm strongly depends on climate and related properties, which reflect the direct or indirect influence of insolation¹⁹. As a result, there is a striking resemblance between δ¹⁸O_atm and mid-June insolation at 65°N for the entire Vostok record (Fig. 3). This provides information on the validity of our glaciological timescale

<figure><img src="image_5.png" /></figure>

**Figure 4** Spectral properties of the Vostok time series. Frequency distribution (in cycles yr⁻¹) of the normalized variance power spectrum (arbitrary units). Spectral analysis was done using the Blackman-Tukey method (calculations were performed with the Analyseries software⁴⁷): **a**, isotopic temperature; **b**, dust; **c**, sodium; **d**, δ¹⁸O_atm; **e**, CO₂; and **f**, CH₄. Vertical lines correspond to periodicities of 100, 41, 23 and 19 kyr.
(see Box 1). The precessional frequencies, which do not account for much variance in $\Delta T_1$, are strongly imprinted in the $\delta^{18}\text{O}_{\text{atm}}$ record (36% of the variance in this band, Fig. 4). In addition, the remarkable agreement observed back to stage 7.5 between the amplitude of the filtered components of the mid-June insolation at 65° N and $\delta^{18}\text{O}_{\text{atm}}$ in the precessional band¹³ holds true over the last four climate cycles (not shown). As suggested by the high variance of $\delta^{18}\text{O}_{\text{atm}}$ in the precessional band, this orbital frequency is also reflected in the Dole effect, the difference between $\delta^{18}\text{O}_{\text{atm}}$ and $\delta^{18}\text{O}_{\text{sw}}$, confirming results obtained on the last two climate cycles¹⁹,³⁰.

**Aerosols.** Figure 2 shows records of aerosols of different origins. The sodium record represents mainly sea-salt aerosol entrained from the ocean surface, whereas the dust record corresponds to the small size fraction (~2 µm) of the aerosol produced by the continent. The extension of the Vostok record confirms much higher fallout during cold glacial periods than during interglacials. Concentrations range up to 120 ng g⁻¹, that is, 3 to 4 times the Holocene value, for sea-salt. For dust, they rise from about 50 ng g⁻¹ during interglacials to 1,000–2,000 ng g⁻¹ during cold stages 2, 4, 6, 8 and 10.

The sodium concentration is closely anti-correlated with isotopic temperature ($r^2 = -0.70$ over the past 420 kyr). The power spectrum of the sodium concentration, like that of $\Delta T_1$, shows periods of around 100, 40 and 20 kyr (Fig. 4). Conditions prevailing during the present-day austral winter could help explain the observed glacial/interglacial changes in sodium. The seasonal increase of marine aerosol observed in the atmosphere and snow at the South Pole in September³¹ corresponds to the maximum extent of sea ice; this is because the more distant source effect is compensated by the greater cyclonic activity, and by the more efficient zonal and meridional atmospheric circulation probably driven by the steeper meridional (ocean–Antarctica) temperature gradient. These modern winter conditions may be an analogue for glacial climates, supporting the apparent close anti-correlation between sodium concentration and temperature at Vostok.

The extension of the Vostok dust record confirms that continental aridity, dust mobilization and transport are more prevalent during glacial climates, as also reflected globally in many dust records (see ref. 10 and references therein). The presence of larger particles in the Vostok record, at least during the Last Glacial Maximum¹⁰, indicates that the atmospheric circulation at high southern latitudes was more turbulent at that time. Lower atmospheric moisture content and reduced hydrological fluxes may also have contributed significantly (that is, one order of magnitude³²) to the very large increases of dust fallout during full glacial periods because of a lower aerosol-removal efficiency.

Unlike sodium concentration, the dust record is not well correlated with temperature (see below) and shows large concentrations of variance in the 100- and 41-kyr spectral bands (Fig. 4). The

Vostok dust record is, in this respect, similar to the tropical Atlantic dust record of de Menocal³³ who attributes these spectral characteristics to the progressive glaciation of the Northern Hemisphere and the greater involvement of the deep ocean circulation. We suggest that there also may be some link between the Vostok dust record and deep ocean circulation through the extension of sea ice in the South Atlantic Ocean, itself thought to be coeval with a reduced deep ocean circulation³⁴. Our suggestion is based on the fact that the dust source for the East Antarctic plateau appears to be South America, most likely the Patagonian plain, during all climate states of the past 420 kyr (refs 35, 36). The extension of sea ice in the South Atlantic during glacial times greatly affects South American climate, with a more northerly position of the polar front and the belt of Westerlies pushed northward over the Andes. This should lead, in these mountainous areas, to intense glacial and fluvial erosion, and to colder and drier climate with extensive dust mobilization (as evidenced by glacial loess deposits in Patagonia). Northward extension of sea ice also leads to a steeper meridional temperature gradient and to more efficient poleward transport. Therefore, Vostok dust peaks would correspond to periods of increased sea-ice extent in the South Atlantic Ocean, probably associated with reduced deep ocean circulation (thus explaining observed similarities with the tropical ocean dust record³³).

**Greenhouse gases.** The extension of the greenhouse-gas record shows that the main trends of CO₂ and CH₄ concentration changes are similar for each glacial cycle (Fig. 3). Major transitions from the lowest to the highest values are associated with glacial–interglacial transitions. At these times, the atmospheric concentrations of CO₂ rises from 180 to 280–300 p.p.m.v. and that of CH₄ rises from 320–350 to 650–770 p.p.b.v. There are significant differences between the CH₄ concentration change associated with deglaciations. Termination III shows the smallest CH₄ increase, whereas termination IV shows the largest (Fig. 5). Differences in the changes over deglaciations are less significant for CO₂. The decrease of CO₂ to the minimum values of glacial times is slower than its increase towards interglacial levels, confirming the sawtooth record of this property. CH₄ also decreases slowly to its background level, but with a series of superimposed peaks whose amplitude decreases during the course of each glaciation. Each CH₄ peak is itself characterized by rapid increases and slower decreases, but our resolution is currently inadequate to capture the detail of millennial-scale CH₄ variations. During glacial inception, Antarctic temperature and CH₄ concentrations decrease in phase. The CO₂ decrease lags the temperature decrease by several kyr and may be either steep (as at the end of interglacials 5.5 and 7.5) or more regular (at the end of interglacials 9.3 and 11.3). The differences in concentration–time profiles of CO₂ and CH₄ are reflected in the power spectra (Fig. 4). The 100-kyr component dominates both CO₂ and CH₄ records. However, the obliquity and precession components are much stronger for CH₄ than for CO₂.

The extension of the greenhouse-gas record shows that present-day levels of CO₂ and CH₄ (~360 p.p.m.v. and ~1,700 p.p.b.v., respectively) are unprecedented during the past 420 kyr. Pre-industrial Holocene levels (~280 p.p.m.v. and ~650 p.p.b.v., respectively) are found during all interglacials, while values higher than these are found in stages 5.5, 9.3 and 11.3 (this last stage is probably incomplete), with the highest values during stage 9.3 (300 p.p.m.v. and 780 p.p.b.v., respectively).

The overall correlation between our CO₂ and CH₄ records and the Antarctic isotopic temperature⁵,⁹,¹⁶ is remarkable ($r^2 = 0.71$ and 0.73 for CO₂ and CH₄, respectively). This high correlation indicates that CO₂ and CH₄ may have contributed to the glacial–interglacial changes over this entire period by amplifying the orbital forcing along with albedo, and possibly other changes¹⁵,¹⁶. We have calculated the direct radiative forcing corresponding to the CO₂, CH₄ and N₂O changes¹⁶. The largest CO₂ change, which occurs between stages 10 and 9, implies a direct radiative warming of

<table><caption>**Table 1 Comparison of the glaciological timescale and orbitally derived information**</caption><thead><tr><th>Depth (m)</th><th>Insolation maximum (kyr)</th><th>Age GT4 (kyr)</th><th>Difference (kyr)</th></tr></thead><tbody><tr><td>305</td><td>11</td><td>10</td><td>1</td></tr><tr><td>900</td><td>58</td><td>57</td><td>1</td></tr><tr><td>1,213</td><td>84</td><td>83</td><td>1</td></tr><tr><td>1,528</td><td>105</td><td>105</td><td>0</td></tr><tr><td>1,863</td><td>128</td><td>128</td><td>0</td></tr><tr><td>2,110</td><td>151</td><td>150</td><td>1</td></tr><tr><td>2,350</td><td>176</td><td>179</td><td>-3</td></tr><tr><td>2,530</td><td>199</td><td>203</td><td>-4</td></tr><tr><td>2,683</td><td>220</td><td>222</td><td>-2</td></tr><tr><td>2,788</td><td>244</td><td>239</td><td>5</td></tr><tr><td>2,863</td><td>265</td><td>255</td><td>10</td></tr><tr><td>2,972</td><td>293</td><td>282</td><td>11</td></tr><tr><td>3,042</td><td>314</td><td>301</td><td>13</td></tr><tr><td>3,119</td><td>335</td><td>322</td><td>13</td></tr></tbody></table>

Control points were derived assuming a correspondence between maximum 65°N mid-June insolation and $\delta^{18}\text{O}_{\text{atm}}$ mid-transitions. Age GT4 refers to the age of the gas obtained after correction for the gas-age/ice-age differences²⁰.
<figure><img src="image_6.png" /><figcaption>**Figure 5** Vostok time series during glacial terminations. Variations with respect to time (GT4) of: **a**, dust; **b**, δD of ice (temperature proxy); **c**, $CO_2$; **d**, $CH_4$; and **e**, $\delta^{18}O_{atm}$</figcaption></figure>

for glacial terminations I to IV and the subsequent interglacial periods (Holocene, stage 5.5, stage 7.5 and stage 9.3).

$\Delta T_{rad} = 0.75^{\circ}C$. Adding the effects of $CH_4$ and $N_2O$ at this termination increases the forcing to $0.95^{\circ}C$ (here we assume that $N_2O$ varies with climate as during termination I$^{37}$). This initial forcing is amplified by positive feedbacks associated with water vapour, sea ice, and possibly clouds (although in a different way for a 'doubled $CO_2$' situation than for a glacial climate$^{38}$). The total glacial-interglacial forcing is important ($\sim 3 \text{ W m}^{-2}$), representing 80% of that corresponding to the difference between a 'doubled $CO_2$' world and modern $CO_2$ climate. Results from various climate simulations$^{39,40}$ make it reasonable to assume that greenhouse gases have, at a global scale, contributed significantly (possibly about half, that is, $2-3^{\circ}C$) to the globally averaged glacial-interglacial temperature change.

## Glacial terminations and interglacials

Our complete Vostok data set allows us to examine all glacial commençements and terminations of the past 420 kyr. We can examine the time course of the following five properties during these events: $\delta D_{ice}$, dust content, $\delta^{18}O_{atm}$, $CO_2$ and $CH_4$ (Fig. 5). We consider that, during the terminations, $\delta^{18}O_{atm}$ tracks $\delta^{18}O_{sw}$ with a lag of $\sim 2$ kyr (ref. 11), the response time of the atmosphere to changes in $\delta^{18}O_{sw}$, $\delta^{18}O_{atm}$ can thus be taken as an indicator of the large ice-volume changes associated with the deglaciations. Broecker and Henderson$^{41}$ recently supported this interpretation for the last two terminations and discussed its limitations. Our extended $\delta^{18}O_{atm}$ record indeed reinforces such an interpretation, as it shows that the amplitudes of $\delta^{18}O_{atm}$ changes parallel $\delta^{18}O_{sw}$ changes for all four terminations. $\delta^{18}O_{sw}$ changes are similar for terminations I, II and IV ($1.1-1.2\%$) but much smaller for termination III ($\sim 0.6\%$). The same is true for $\delta^{18}O_{atm}$ ($1.4-1.5\%$ for I, II and IV, and $0.8\%$ for III).

A striking feature of the Vostok deuterium record is that the Holocene, which has already lasted 11 kyr, is, by far, the longest stable warm period recorded in Antarctica during the past 420 kyr (Fig. 5). Interglacials 5.5 and 9.3 are different from the Holocene, but similar to each other in duration, shape and amplitude. During each of these two events, there is a warm period of $\sim 4$ kyr followed by a relatively rapid cooling and then a slower temperature decrease (Fig. 5), rather like some North Atlantic deep-sea core records$^{42}$. Stage 7.5 is different in all respects, with a slightly colder maximum,

a more spiky shape, and a much shorter duration (7 kyr at mid-transition compared with 17 and 20 kyr for stages 5.5 and 9.3, respectively). This difference between stage 7.5 and stages 5.5 or 9.3 may result from the different configuration of the Earth's orbit (in particular concerning the phase of precession with respect to obliquity$^3$). Termination III is also peculiar as far as terrestrial aerosol fallout is concerned. Terminations I, II and IV are marked by a large decrease in dust; high glacial values drop to low interglacial values by the mid-point of the $\delta D_{ice}$ increase. But for termination III, the dust concentration decreases much earlier, with low interglacial values obtained just before a slight cooling event, as for termination I (for which interglacial values are reached just before the 'Antarctic Cold Reversal').

Unlike termination I, other terminations show, with our present resolution, no clear temperature anomalies equivalent to the Antarctic Cold Reversal$^6$ (except possibly at the very beginning of termination III). There are also no older counterparts to the Younger Dryas $CH_4$ minimum$^{43}$) during terminations II, III and IV given the present resolution of the $CH_4$ record (which is no better than 1,000–2,000 yr before stage 5).

The sequence of events during terminations III and IV is the same as that previously observed for terminations I and II. Vostok temperature, $CO_2$ and $CH_4$ increase in phase during terminations. Uncertainty in the phasing comes mainly from the sampling frequency and the ubiquitous uncertainty in gas-age/ice-age differences (which are well over $\pm 1$ kyr during glaciations and terminations). In a recent paper, Fischer et al.$^{44}$ present a $CO_2$ record, from Vostok core, spanning the past three glacial terminations. They conclude that $CO_2$ concentration increases lagged Antarctic warmings by 600 $\pm$ 400 years. However, considering the large gas-age/ice-age uncertainty (1,000 years, or even more if we consider the accumulation-rate uncertainty), we feel that it is premature to infer the sign of the phase relationship between $CO_2$ and temperature at the start of terminations. We also note that their discussion relates to early deglacial changes, not the entire transitions.

An intriguing aspect of the deglacial $CH_4$ curves is that the atmospheric concentration of $CH_4$ rises slowly, then jumps to a maximum value during the last half of the deglacial temperature rise. For termination I, the $CH_4$ jump corresponds to a rapid Northern Hemisphere warming (Bölling/Allerød) and an increase
in the rate of Northern Hemisphere deglaciation (meltwater pulse IA)⁴³. We speculate that the same is true for terminations II, III and IV. Supportive evidence comes from the δ¹⁸O_atm curves. During each termination, δ¹⁸O_atm begins falling rapidly, signalling intense deglaciation, within 1 kyr of the CH₄ jump. The lag of deglaciation and Northern Hemisphere warming with respect to Vostok temperature warming is apparently greater during terminations II and IV (~9 kyr) than during terminations I and III (~4–6 kyr). The changes in northern summer insolation maxima are higher during terminations II and IV, whereas the preceding southern summer insolation maxima are higher during terminations I and III. We speculate that variability in phasing from one termination to the next reflects differences in insolation curves⁴¹ or patterns of abyssal circulation during glacial maximum. Our results suggest that the same sequence of climate forcings occurred during each termination: orbital forcing (possibly through local insolation changes, but this is speculative as we have poor absolute dating) followed by two strong amplifiers, with greenhouse gases acting first, and then deglaciation enhancement via ice-albedo feedback. The end of the deglaciation is then characterized by a clear CO₂ maximum for terminations II, III and IV, while this feature is less marked for the Holocene.

Comparison of CO₂ atmospheric concentration changes with variations of other properties illuminates oceanic processes influencing glacial-interglacial CO₂ changes. As already noted for terminations I and II⁴¹, the sequence of climate events described above rules out the possibility that rising sea level induces the CO₂ increase at the beginning of terminations. On the other hand, the small CO₂ variations associated with Heinrich events⁴⁵ suggest that the formation of North Atlantic Deep Water does not have a large effect on CO₂ concentrations. Our record shows similar relative amplitudes of atmospheric CO₂ and Vostok temperature changes for the four terminations. Also, values of both CO₂ and temperature are significantly higher during stage 7.5 than during stages 7.1 and 7.3, whereas the deep-sea core ice volume record exhibits similar levels for these three stages. These similarities between changes in atmospheric CO₂ and Antarctic temperature suggest that the oceanic area around Antarctica plays a role in the long-term CO₂ change. An influence of high southern latitudes is also suggested by the comparison with the dust profile, which exhibits a maximum during the periods of lowest CO₂. The link between dust and CO₂ variations could be through the atmospheric input of iron⁴⁶. Alternatively, we suggest a link through deep ocean circulation and sea ice extent in the Southern Ocean, both of which play a role in ocean CO₂ ventilation and, as suggested above, in the dust input over East Antarctica.

## New constraints on past climate change

As judged from Vostok records, climate has almost always been in a state of change during the past 420 kyr but within stable bounds (that is, there are maximum and minimum values of climate properties between which climate oscillates). Significant features of the most recent glacial-interglacial cycle are observed in earlier cycles. Spectral analysis emphasises the dominance of the 100-kyr cycle for all six data series except δ¹⁸O_atm and a strong imprint of 40- and/or 20-kyr periodicities despite the fact that the glaciological dating is tuned by fitting only two control points in the 100-kyr band.

Properties change in the following sequence during each of the last four glacial terminations, as recorded in Vostok. First, the temperature and atmospheric concentrations of CO₂ and CH₄ rise steadily, whereas the dust input decreases. During the last half of the temperature rise, there is a rapid increase in CH₄. This event coincides with the start of the δ¹⁸O_atm decrease. We believe that the rapid CH₄ rise also signifies warming in Greenland, and that the deglacial δ¹⁸O_atm decrease records rapid melting of the Northern Hemisphere ice sheets. These results suggest that the same sequence

of climate forcing operated during each termination: orbital forcing (with a possible contribution of local insolation changes) followed by two strong amplifiers, greenhouse gases acting first, then deglaciation and ice-albedo feedback. Our data suggest a significant role of the Southern Ocean in regulating the long-term changes of atmospheric CO₂.

The Antarctic temperature was warmer, and atmospheric CO₂ and CH₄ concentrations were higher, during interglacials 5.5 and 9.3 than during the Holocene and interglacial 7.5. The temporal evolution and duration of stages 5.5 and 9.3 are indeed remarkably similar for all properties recorded in Vostok ice and entrapped gases. As judged from the Vostok record, the long, stable Holocene is a unique feature of climate during the past 420 kyr, with possibly profound implications for evolution and the development of civilizations. Finally, CO₂ and CH₄ concentrations are strongly correlated with Antarctic temperatures; this is because, overall, our results support the idea that greenhouse gases have contributed significantly to the glacial-interglacial change. This correlation, together with the uniquely elevated concentrations of these gases today, is of relevance with respect to the continuing debate on the future of Earth's climate.

Received 20 January; accepted 14 April 1999.

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**Supplementary information** is available on *Nature's* World-Wide Web site (http://www.nature.com) or as paper copy from the London editorial office of *Nature*.

**Acknowledgements.** This work is part of a joint project between Russia, France and USA. We thank the drillers from the St Petersburg Mining Institute; the Russian, French and US participants for field work and ice sampling; and the Russian Antarctic Expeditions (RAE), the Institut Français de Recherches et Technologies Polaires (IFRTP) and the Division of Polar Programs (NSF) for the logistic support. The project is supported in Russia by the Russian Ministry of Sciences, in France by PNEDC (Programme National d'Etudes de la Dynamique du Climat), by Fondation de France and by the CEC (Commission of European Communities) Environment Programme, and in the US by the NSF Science Foundation.

Correspondence and requests for materials should be addressed to J.R.P. (e-mail: petit@glaciog.ujf-grenoble.fr).