Trends, Rhythms, and Aberrations in Global Climate 65 Ma to Present James Zachos, Mark Pagani, Lisa Sloan, Ellen Thomas, and Katharina Billups

Supplementary Material

Web note 1. Chronology: Most of the pre-Pliocene deep-sea records published prior to ~1992 were calibrated to now obsolete versions of the Geomagnetic Polarity Time Scale (GPTS). As a result, to assemble a global deep-sea isotope record, it was necessary to first revise and update the age models of each record relative to the GPTS (11). In the process of updating chronologies, we discovered that the biozonation schemes for some sites were either outdated or unreliable due to low sampling resolution or a lack of detailed assemblage information. In those cases, biostratigraphic interpretation was augmented with isotope stratigraphy. The site-to-site correlations are internally consistent ensuring that the primary results will not be affected by future refinements of the GPTS. The Plio-Pleistocene ages are constrained by oxygen isotope records tuned to Northern Hemisphere summer insolation at 65°N. In this approach, the D¹⁸O record is filtered at the appropriate orbital band and correlated to the orbital parameter assuming a constant lag between the insolation forcing and climate response (e.g., ice-volume). The Plio-Pleistocene record have been tuned to the astronomical solutions of Berger and Loutre (12). Only one pre-Pliocene record is orbitally tuned, Site 929 (13). In this case, the astronomical calibration was achieved by aligning lithologic cycles as reflected in magnetic susceptibility records with orbital cycles in a target curve derived from the equations of Laskar (14).

Web note 2. A Primer on Stable Isotope Proxies. Temperature influences the degree of ¹⁸O/¹⁶O fractionation between seawater and the carbonate ions (CO₃) that are incorporated into the shell lattice. For certain taxa, particularly in the bottom dwelling benthic foraminifera, the temperature-fractionation relationship as quantified by field and laboratory calibration (1) appears to remain constant with time, thus allowing for estimation of past water temperature from fossil shell D¹⁸O values. Ice-sheet growth/decay also influences the ¹⁸O/¹⁶O of shell calcite as the mean D¹⁸O of water that evaporates from the ocean and accumulates in ice-sheets is significantly lower (-15 to -40‰ vSMOW) than the D¹⁸O of mean ocean water (0.0‰ vSMOW). Consequently, growth of ice-sheets increases ocean D¹⁸O, by as much as 1.2‰ with full-scale Antarctic ice-sheets, and by an additional ~1.2‰ when N. Hemisphere ice-sheets are included. This dual control on shell D¹⁸O complicates the interpretation of trends over those intervals when large ice-sheets are present. Many strategies have been devised to separate the contributions of ice-volume and temperature, though all have limitations (2, 3, 4, 5, 6).

The stable carbon isotope ratio (¹³C/¹²C) of benthic foraminifera reflects largely on the isotope ratio of the total dissolved inorganic carbon (DIC) of ambient seawater. As such, changes in the D¹³C_DIC of seawater are reflected in the D¹³C of fossil shells. The mean D¹³C_DIC of ocean water (~0.0‰) varies in response to fluctuations in the major fluxes of global carbon cycle (7). For example, algal organic matter is isotopically light (~-20‰) due to the preferential uptake of ¹²C during photosynthesis. Thus, if the global average rate of organic carbon removal (i.e., burial) increases for a sustained period (>10⁴y), it should drive the D¹³C of the remaining DIC higher, assuming other fluxes remain fixed. Similarly, sustained changes in the composition and/or flux of carbon into the ocean from continental weathering and volcanic outgassing can alter mean ocean D¹³C_DIC. On a local scale, the D¹³C_DIC of deep-waters decreases with distance from the source with the addition of ¹²C enriched CO₂ from oxidation of organic matter (8). Consequently, when physical barrier arise to restrict or cutoff deepwater circulation between basins, detectable gradients in D¹³C_DIC may exist and be used to distinguish deep-water masses (9, 10).

Web note 3. Sampling Biases: One of the limitations on reconstructing long-term secular variations is the highly uneven distribution of deep-sea stable isotope data in both space and in time. The global signal for some key intervals is based on data from just a few records. In general, these spatial biases increase with age, moving toward the Atlantic, and shallower water depths. In other words, the Pacific, and abyssal portions of the oceans tend to be under-represented in existing stable isotope records. These biases do not pose a problem for our temperature/ice-volume reconstruction of the late Neogene oceans which were thermally homogeneous. Such biases, however, are a concern for establishing the mean climate-state of some "warmer" time intervals when the thermal gradients within the deep-sea were greater. As for the global carbon isotope record, spatial biases are more important toward the younger part of the record because of circulation related basin to basin fractionation which results in a more than a 1.0‰ difference between the deep Pacific and Atlantic. Isotope records indicate that this inter-basin gradient developed sometime in the mid-late Miocene, and intensified in the Pliocene (10, 15, 16). Prior to the late Miocene, because of the lack of low-latitude barriers to deep communication and weak to non-existent NADW production, the carbon isotope chemistry of the deep basins was more homogeneous. Additional bias is introduced into the compilation by differences in sampling density in each record. In general, the majority of stable isotope records generated in the last decade are of a much higher resolution than older records (pre-1992). This bias applies mainly to the Paleogene interval for which there are only a few high-resolution (10³-10⁴y) records (e.g., Sites 522, 744, 689, and 929).

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Supplemental Table 1. Source of isotope records used to construct figure 1.
Site	Ocean	Genus	Source
77	PAC	CIB	(17)
206	PAC	CIB	(18)
213	IND	NUT	(5)
215	IND	NUT	(5)
360	ATL	CIB	(19)
384	ATL	Gavelinella + NUTT	(20)
522	ATL	CIB	(21)
523	ATL	NUTT	(22)
526	ATL	CIB	(23)
527	ATL	NUTT	(23)
553	ATL	CIB	(24)
558	ATL	CIB	(25)
563	ATL	CIB	(16)
574	PAC	CIB	(15)
577	PAC	NUTT	(26)
577	PAC	NUTT	(27)
588	PAC	CIB	(28)
590	PAC	CIB	(18)
591	PAC	CIB	(18)
593	PAC	CIB	(18)
594	PAC	CIB	(18)
607	ATL	CIB	(29)
608	ATL	CIB	(16)
659	ATL	CIB	(30)
689	ATL	CIB & NUTT	(31, 32)
690	ATL	CIB & NUTT	(31, 33)
704	ATL	CIB	(34)
738	IND	CIB & NUTT	(35)
744	IND	CIB	(21)
747	IND	CIB	(36)
748	IND	CIB	(37)
752	IND	CIB	(38)
754	IND	Mixed	(39)
757	IND	CIB	(5)
758	IND	NUTT	(38)
806	PAC	CIB	(40)
846	PAC	CIB	(41)
849	PAC	CIB	(42)
865	PAC	CIB & NUTT	(43)
926	ATL	mixed	(44, 45)
929	ATL	CIB	(46, 47, 48)
959	ATL	CIB	(49)
CIB - Cibicidoides
NUTT - Nuttallides

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