5.6.2. The “B-A Ventilation Surge”

The “B-A ventilation surge” coincides with what has been interpreted to reflect an intensification (Gherardi et al., 2005; McManus et al., 2004), and possibly also an “extreme deepening” or “overshoot,” of the AMOC (Barker et al., 2010; Skinner et al., 2013). A clear association is thus apparent between the rapid rejuvenation of the deep (and shallow sub-surface) North Atlantic, the reinvigoration of the AMOC, and the abrupt warming over Greenland that marks the onset of the B-A, and that coincides with warming and sea-ice loss in the North Atlantic region. This is entirely consistent with the standard account of the AMOC's role in the thermal bi-polar seesaw (Barker et al., 2009). However, it is intriguing that the bulk of the collected B-Atm records from the deep North Atlantic, shown in Figure 13e, appear to suggest that rejuvenation may have preceded the abrupt B-A onset by several centuries. While this may simply reflect signal smoothing due to bioturbation, or limitations in chronology precision, it is intriguing that a similar lead has been substantiated for the analogous transition that occurred at the end of the YD (Muschitiello et al., 2019).

While the influence of an AMOC during the B-A surge might be expected to reach downstream into the deep-
and intermediate South Atlantic within centuries (Barker et al., 2010; Skinner et al., 2010), the time required to
advect this signal to the deep EEP, and subsequently the deep North Pacific, would be >1000 years (Primeau &
Deleersnijder, 2009). Nevertheless, it has been proposed by Umling and Thunell (2017) that a pulse of radiocar-
bon ventilation at ~2,700 m in the deep EEP was essentially synchronous with the B-A surge observed in the
North Atlantic at ~14.8 ka BP, yet was delayed with respect to the earlier rejuvenation initiated in the Southern
Ocean and intermediate Atlantic (which represents an interface between northern and southern sourced waters
[Freeman et al., 2015]). The deep EEP B-Atm ventilation pulse inferred by Umling and Thunell (2017) at ~14.8
ka BP is not clearly expressed in B-P ¹⁴C-age offsets at this site. As concluded by Umling and Thunell (2017), this
would require that the radiocarbon ventilation of the thermocline- and deep EEP were affected in synchrony by a
presumed southern source. This could be achieved in principle by rapid changes in Southern Ocean surface reser-
voir ages at ~14.8 ka BP (i.e., air-sea gas exchange, influencing Antarctic mode- and deep waters transported to
the EEP thermocline) (de la Fuente et al., 2015; Skinner et al., 2015). While Southern Ocean surface reservoir
ages appear not to support this (exhibiting a gradual decline across HS1, in asynchrony with the North Atlantic)
(Skinner, Muschitiello, & Scrivner, 2019) (Figures 13c and 14c), an abrupt centennial ventilation pulse ~14.8
ka BP does seem to have affected the shallow sub-surface of the Drake Passage (Li et al., 2020) and Galapagos
(Chen et al., 2020) (see below). Alternatively, the North Pacific may have contributed, alongside the AMOC, to a
surge in deep Pacific ventilation during the B-A (Gray et al., 2018). There is a great deal of scatter in the collected
data, though there does appear to be a slight decrease in ventilation ages, on average, just prior to the B-A in the
deep (>2 km) North Pacific. Nevertheless, this trend appears to be reversed during the B-A, in keeping with an
anti-phased relationship between PMOC and AMOC intensity (Chikamoto, Abe-Ouchi, et al., 2012; Chikamoto,
Menviel, et al., 2012; Max et al., 2014; Menviel et al., 2014; Mix et al., 1999; Okazaki et al., 2010, 2012). In light
of the foregoing discussion, it is fair to say that a host of interesting and puzzling observations from the Pacific
basin currently defy a tidy explanation, particular in relation to the B-A. Their elucidation represents an important
target for future work, particularly as this bears on the distinct roles of convection and air-sea exchange in the high
latitudes of each hemisphere.

A further puzzle emerges when considering that the AMOC surge across the B-A appears to have had little impact
on atmospheric CO2, whereas an acceleration of ocean ventilation via renewed deep convection would be expected
to reduce the ocean's respired carbon content and push up atmospheric CO2 (Sarmiento & Toggweiler, 1984).
Instead, CO2 declined slightly across the B-A (Marcott et al., 2014), though notably far more slowly than during
previous D-O interstadials of the last glacial (Wolff et al., 2009). Indeed, D-O interstadials of the last glacial
period (which were all characterized by a relatively strong AMOC [Gottschalk et al., 2015; Henry et al., 2016])
were consistently associated with declining atmospheric CO2 (Ahn & Brook, 2008; Bauska et al., 2021), unlike
during the B-A. The nearly stable CO2 levels across the B-A therefore suggest an additional source of carbon (or
a missing carbon sink) as compared to preceding D-O interstadials (Skinner et al., 2013, 2021).

Part of the explanation for nearly stable atmospheric CO₂ during the BA may be that the AMOC has a limited
ability to affect the ocean's total respired carbon content (Hain et al., 2014), both due to its limited volumetric
influence (Johnson, 2008; Primeau, 2005), and relative to the overwhelming impacts of countervailing changes
in ventilation occurring in the Southern Ocean and North Pacific during D-O interstadials (Anderson et al., 2009;
Freeman et al., 2015; Menviel et al., 2014, 2015; Okazaki et al., 2012; Skinner et al., 2020). Therefore, from a
marine CO₂ release perspective, the B-A may have simply been an intermission between stronger pulses of CO₂ release during HS1 and the YD, while from an atmospheric radiocarbon perspective, the onset of the B-A may have been the more significant event (Hain et al., 2014). In this respect it is worth noting that the net fluxes of radiocarbon and CO₂ across the air-sea interface do not have to be correlated (see Section 3.1 above). Taken together, these considerations may suggest that deglacial AMOC changes associated with the “thermal bipolar seesaw” primarily influenced atmospheric CO₂ through far-field impacts during periods of reduced AMOC (Cheng et al., 2009), in particular in the Southern Ocean (Anderson et al., 2009; Skinner et al., 2014), and perhaps also the North Pacific (Menviel et al., 2014; Okazaki et al., 2010) (Figure 14). Accordingly, pulses of ocean interior ventilation during HS1 and the YD might be seen as the main enablers for deglacial CO₂ rise: metaphorically, “punctures” that produced a leak in the biological pump and thus deflated the ocean's overall sequestered carbon inventory (Anderson et al., 2009; Marchitto et al., 2007). While this canonical view of the AMOC contribution to deglacial carbon cycle dynamics is supported by many available observations and model simulations (e.g., Anderson et al., 2009; Cheng et al., 2009; Marchitto et al., 2007; Menviel et al., 2014, 2018; Okazaki et al., 2010; Skinner et al., 2014), it is now clear that it can only represent a partial explanation for deglacial CO₂ change as it ignores the possible contribution of shorter lived centennial events that occurred withinstadial periods and at stadial-interstadial transitions (Bauska et al., 2021; Marcott et al., 2014). These are taken up next.

5.6.3. Centennial “Jumps”

Highly resolved CO₂ reconstructions from the WAIS Divide ice-core demonstrate that millennial-scale perturbations linked to the thermal bipolar seesaw can only account for a portion of deglacial CO₂ rise: at least ~40% of the total deglacial rise was achieved during separate centennial scale CO₂ jumps that were far more short-lived than HS1 and the YD (Marcott et al., 2014). Three main abrupt CO₂ jumps have been identified: in the middle of HS1 (16.3 ka BP), at the end of HS1 (and the onset of the B-A; 14.8 ka BP), and at the end of the YD (11.7 ka BP) (Marcott et al., 2014). Although evidence for marine radiocarbon ventilation pulses coinciding with each of these CO₂ jumps is hard to identify in the bulk of the available marine radiocarbon data (mostly foraminifer-based; see Figure 13), some tentative evidence for an association between ocean ventilation and centennial atmospheric CO₂ (and atmospheric Δ¹⁴C) jumps does exist. This is shown in Figure 15d, where rather subtle pulses in radiocarbon ventilation are apparent at ~12, 15, and 16 ka BP in collected coral-based data from both the Drake Passage in the Southern Ocean (in particular <1 km water depth) (Burke & Robinson, 2012; Chen et al., 2015; Li et al., 2020), and the Galapagos Plateau (~600 m water depth) (Chen et al., 2020). These regions are plausibly linked hydrographically by the influence of southern sourced intermediate-depth mode water.

The coincidence of the subtle ventilation pulses shown in Figure 15d with periods of rapid atmospheric CO₂ rise (Figure 15b) is probably not spurious, due to the accuracy of U-series dating performed on the corals. However, it is likely that the expression of ventilation pulses that coincided with the atmospheric CO₂ jumps may have been obscured to some degree (particularly in the Drake Passage data) due to the scatter of coral samples in space and time, which precludes the observation of varying conditions at a single location over time. While foraminifer-based radiocarbon ventilation data have the advantage of providing consistently ordered observations at a single location over time (in the absence of sedimentary disturbance), they are more difficult to date with centennial precision, and are typically affected by the smoothing and potentially biasing effects of bioturbation (see Section 2.2 above). Nevertheless, a few foraminifer-based radiocarbon ventilation reconstructions, for example, from the Brazil Margin (similarly influenced by southern-sourced water to some degree), provide rather tentative circumstantial evidence for declining B-Atm ventilation ages ~12, 15, and 16ka BP (see Figure 15e).

Two of the CO₂ jumps (at the end of HS1 and the YD) coincide with well-documented intervals of rapid intensification of the AMOC (McManus et al., 2004), and associated radiocarbon ventilation via the intermediate- (Chen et al., 2015) and deep Atlantic (Robinson et al., 2005; Skinner et al., 2014, 2021; Figures 13d and 13e). A link with North Atlantic convection is also suggested by rapid radiocarbon ventilation pulses that have been reconstructed in the North Atlantic south of Iceland (Thornalley et al., 2011). Tentatively therefore, the ventilation pulses seen in the Southern Ocean, and downstream in the equatorial Pacific (see Figure 15), may have originated in the North Atlantic. Indeed, it has been proposed, on the basis of pH reconstructions (Rae et al., 2018), that the CO₂ jumps at ~11.7 and 14.8 ka BP may have been linked to carbon loss from the Southern Ocean due to the “excavation” of respired carbon from the deep Atlantic basin by rapid AMOC intensification at the onset of interstadial conditions in the North Atlantic, which was made possible by Antarctic sea-ice loss during the preceding “swing” of the thermal bipolar seesaw that warmed the South Atlantic (Rae et al., 2018).
<figure><img src="image_20.png" /></figure>

**Figure 15.** Illustration of centennial-scale CO₂ changes and associated marine radiocarbon ventilation anomalies. (a) Time-derivative of 600-year moving average of Intcal20 atmospheric Δ¹⁴C_atm (Reimer et al., 2020). (b) time-derivative of 500-year moving average of WAIS Divide atmospheric CO₂ (Marcott et al., 2014). (c) Atmospheric δ¹³CO₂, based on: compiled measurements from a suite of Antarctic ice cores (Schmitt et al., 2012) (light blue filled circles and cubic spline); and based on measurements in Taylor Glacier ice (Bauska et al., 2016) (black filled squares and line). (d) Coral radiocarbon ventilation (B-Atm) data: from the Drake Passage (filled blue circles are data >1,500 m water depth; orange filled circles are data <1,500 m water depth; solid and dashed red lines show cubic spline through all data, with 95% confidence range; Burke & Robinson, 2012; Chen et al., 2015; Li et al., 2020); from 627 m water depth off Galapagos (black filled circles and line, Chen et al., 2020). (e) Foraminifer-based B-Atm data from 1,000 m (filled gray circles and line) and 3,400 m (filled blue circles and line), on the Brazil Margin (Skinner et al., 2021). Vertical shaded bars indicate the approximate timing of HS1 and the YD, with darker bars indicating the timing of three centennial CO₂ pulses. All ventilation ages have been “corrected” by the difference between the Intcal20 and Intcal13 atmospheric radiocarbon curves.

The temporal resolution of measurements in the latter study (Rae et al., 2018) was too low to rule in or out a similar dynamic for the third CO₂ jump at ~16.3 ka BP, and other evidence has previously been advanced that would tentatively link the mid-HS1 CO₂ pulse to a millennial ventilation anomaly in the North Pacific instead (Rae et al., 2014). However, if it is hypothesized that all three of the main centennial CO₂ jumps during the last deglaciation were linked to the same mechanism, it is not obvious that a reinvigoration of the AMOC was at the center of it, mainly because a mid-HS1 AMOC pulse has so far not been identified (McManus et al., 2004). Nevertheless, the emerging radiocarbon evidence for a mid-HS1 ventilation pulse in the South Atlantic and equatorial Pacific (Figure 15), and perhaps also the North Atlantic (Thornalley et al., 2011), suggests that a marine role in the deglacial centennial CO₂ jumps should not yet be ruled out. Indeed, even if a mid-HS1 AMOC surge has not yet been identified, a change in the geometry of the AMOC is apparent within HS1, when comparing Pa/Th
reconstructions of ocean transports in the western and eastern North Atlantic basins (Gherardi et al., 2005; Ng et al., 2018).

The existence of two distinct phases during the HS1 chronozone has long been apparent in a range of archives, including: atmospheric CO₂ (Marcott et al., 2014; Monnin et al., 2001; Figure 15b); atmospheric δ¹³CO₂ (Bauska et al., 2016; Figure 15c); atmospheric methane (Rhodes et al., 2015); low latitude hydrological cycling (Wang et al., 2001); North American Great Plains hydrology (Broecker et al., 2009); ice-rafted debris deposition in the North Atlantic (Bard et al., 2000); and surface temperature variability on the Iberian Margin (Pailler & Bard, 2002; Skinner & Elderfield, 2005; Skinner & Shackleton, 2006). Piecing these lines of evidence together into a coherent mechanistic framework remains a difficult challenge to address (Broecker & Putnam, 2012), and is one that appears to be intimately connected to an elucidation of the mechanisms behind the centennial CO₂ jumps of the last deglaciation (Bauska et al., 2021).

Despite the plausibility of an ocean ventilation role in at least some of the deglacial centennial CO₂ jumps (Bauska et al., 2016, 2021), and the emerging support for this role from some observations (Figure 15; Chen et al., 2015, 2020; Li et al., 2020; Rae et al., 2018), other explanations for these rapid carbon cycle perturbations have also been advanced. It has been proposed for example, that the atmospheric CO₂ pulse at ~14.8 ka BP may have resulted from northern hemisphere permafrost thawing, possibly as a knock-on effect of abrupt AMOC intensification and North Atlantic warming (Köhler et al., 2014). Indeed, it is possible that each of the centennial pulses involved the activation of different carbon sources, of either terrestrial or marine origin (Bauska et al., 2016; Chen et al., 2015; Köhler et al., 2014; Winterfeld et al., 2018), leaving open the possibility that (regardless of an ocean ventilation contribution) the correct mix of sources for all three of the deglacial pulses has yet to be identified.

Independent of their explanation, these abrupt CO₂ jumps, and the slower millennial increases across HS1 and the YD, emphasize the important role of rapid adjustments in the deglacial processes: deglacial carbon cycle evolution was clearly by “by jerks, *and* by creeps.” An important next step, for our understanding of the process of deglaciation and the involvement of the carbon cycle, will surely be to identify the origins of (and possible synergies between) events that occurred on a range of timescales: abrupt centennial jumps that were dispersed throughout the deglacial period; millennial anomalies that were coordinated by the bipolar seesaw; and longer term “mode shifts” (which may have been completed already half-way through the deglaciation).

### 5.7. Extreme Deglacial Anomalies: Fly in the Ointment or Red Herring

As noted above, there is a final grouping of deglacial radiocarbon ventilation records that resists integration into the canonical narrative of ocean circulation change and deglacial marine carbon release to the atmosphere, or indeed the general framework described above. These records derive primarily from the intermediate-depth eastern low-latitude Pacific (Baja California, California Margin, Galapagos), and exhibit a diversity of deglacial signatures. Some, from Baja California, exhibit well-expressed transient millennial-scale radiocarbon depletions that appear to coincide with the North Atlantic stadials HS1 and the YD (Lindsay et al., 2016; Marchitto et al., 2007). Others, mainly from just outside of the Gulf of California, exhibit a broad (and rather noisy) “extreme” radiocarbon anomaly that straddles the entire deglacial period, from the onset of HS1 to the end of the Pre-boreal period ~10 ka BP (Rafter et al., 2018, 2019). Yet others, from Galapagos, exhibit an extreme radiocarbon depletion that developed well before the LGM and was eliminated by the onset of the Holocene, ~10 ka BP (Bova et al., 2018; Stott et al., 2009). Similarly large amplitude radiocarbon depletions have also been reported at two sites in the South Pacific (Ronge et al., 2016), and a site in the Arabian Sea (Bryan et al., 2010) (the latter bearing a passing resemblance to the signal observed off Baja California) (Lindsay et al., 2016; Marchitto et al., 2007).

Suspicions that these observations reflect diagenetic overprinting, bioturbation, turbidites, or inaccurate age-models, have been argued against on the basis of the apparent consistency between sites within specific regions (Bova et al., 2018; Lindsay et al., 2016; Marchitto et al., 2007; Rafter et al., 2019; Stott et al., 2009), and due to a lack of evidence for sedimentary disturbance, secondary calcite phases, or age-differences between foraminifera of different species or of different degrees of preservation (Rafter et al., 2019; Stott et al., 2009). Nevertheless, high resolution radiocarbon dating has shown that at least one of the two South Pacific sites that exhibit “extreme” radiocarbon depletion (Ronge et al., 2016) has indeed been affected by a sedimentation hiatus and bioturbation (Ronge et al., 2019). Depositional hiatuses and age-reversals are also apparent in a set of cores from the
<figure><img src="image_21.png" /></figure>

**Figure 16.** Southern Ocean radiocarbon ventilation data, south of 40°S and >2 km water depth. (a) Comparison of southwest Pacific B-Atm offsets from core PS75/100-4 (gray crossed circles and solid line, Ronge et al., 2016), B-Atm associated with tephra deposited off New Zealand (black open stars [Sikes et al., 2000], and black filled stars [Sikes, Cook, & Guilderson, 2016]), and spline fit to Southern Ocean data >2 km water depth shown in plots B to D (solid blue line and 95% shaded confidence interval). (b) Deep South Pacific and EEP B-Atm offsets (Dai et al., 2021; de la Fuente et al., 2015; Ronge et al., 2016; Skinner et al., 2015; Umling & Thunell, 2017). (c) Deep South Indian B-Atm offsets (Gottschalk et al., 2020; Ronge et al., 2020). (d) Deep South Atlantic B-Atm offsets (Barker et al., 2010; Diz & Barker, 2015; Skinner et al., 2010). All ventilation ages have been “corrected” according to the difference between the Intcal20 and Intcal13 atmospheric radiocarbon curves.

southern Indian Ocean presented by Ronge et al. (2020), further underlining the challenges that can arise (Costa et al., 2017), and that can often go undetected in some sedimentary contexts (Stott, 2020).

It is worth noting that additional observations of B-Atm offsets ranging from ~1,600 up to 6,800 ¹⁴C years have also been reconstructed off New Zealand in association with the Kawakawa tephra, at 25,650 ± 40 kaBP (Sikes et al., 2000; Sikes, Cook, & Guilderson, 2016) (see Figure 16a). These results indicate a consistent “aging” of deep waters in the South Pacific during the last glacial period, consistent with numerous other reconstructions from the across the South Pacific water column (Hines et al., 2015; Shao et al., 2019; Siani et al., 2013), as well as from the Atlantic (Barker et al., 2010; Skinner et al., 2010), Indian (Gottschalk et al., 2020; Ronge et al., 2020),
and Pacific sectors of the deep (>2 km) Southern Ocean (Dai et al., 2021; Skinner et al., 2015; see Figure 16). However, the large degree of scatter exhibited in the results obtained for a single time horizon also suggests that individual observations (which may be subject to bioturbational biases above/below tephra [Sikes, Cook, & Guilderson, 2016]) are probably less representative of regional ventilation ages than their average. On this basis, only a single record from south of New Zealand (Ronge et al., 2016) would appear to deviate from a significant body of fairly consistent evidence from across the deep Southern Ocean (Barker et al., 2010; Dai et al., 2021; Gottschalk et al., 2020; Shao et al., 2019; Sikes, Cook, & Guilderson, 2016; Skinner et al., 2010, 2015, 2017), and indeed the deep EEP (de la Fuente et al., 2015; Umling & Thunell, 2017) (see Figure 16). It is probably fair to say that the current balance of evidence does not support the widespread prevalence of extremely aged *deep-water* radiocarbon signatures (i.e., >4,000 <sup>14</sup>Cyrs equivalent) in the South Pacific at the LGM or during deglaciation.

Off Baja California (see Figure 17d), marked radiocarbon anomalies during HS1 and the YD have been reproduced from a range of intermediate water depths, 430–1,270 m (Lindsay et al., 2016; Marchitto et al., 2007), and these are tracked by similar anomalies in obligate surface dwelling (i.e., algal symbiont bearing) planktonic foraminifera (red/orange symbols and lines in Figure 17d). The amplitude of the planktonic radiocarbon anomalies is ~50% of the benthic anomalies, and has been interpreted to reflect upwelling of “aged” intermediate waters sourced from the Southern Ocean (Lindsay et al., 2015). This interpretation has drawn circumstantial support from surface water pCO₂ reconstructions (Martinez-Boti et al., 2015), and some surface water “reservoir age” reconstructions from the EEP (de la Fuente et al., 2015; Umling & Thunell, 2017). However, all of these studies appear to be inconsistent with surface reservoir ages derived from radiocarbon dating of allochthonous wood detritus in a marine core from the Columbian Margin (Zhao & Keigwin, 2018).

In the modern ocean, intermediate waters off Baja California are primarily sourced from the North Pacific, with a thin wedge of Antarctic Intermediate Water (AAIW) beneath this (Fiedler & Talley, 2006). Although a North Pacific origin for extreme radiocarbon anomalies in the shallow eastern low-latitude Pacific has been proposed (Bova et al., 2018), radiocarbon ventilation ages in the deep North Pacific were apparently never >2,500 <sup>14</sup>Cyrs since the LGM (Galbraith et al., 2007; Lund, Adkins, & Ferrari, 2011; Lund, Mix, & Southon, 2011; Mix et al., 1999; Okazaki et al., 2010, 2012; Rae et al., 2014) (see Figure 17e), making it difficult to account for the observed radiocarbon anomalies off Baja California with a northern source. Moreover, a southern source is also problematic, as intermediate depth sites in the Southern Pacific and Drake Passage do not record evidence for AAIW as old as implied by the Baja California anomalies (Burke & Robinson, 2012; De Pol-Holz et al., 2010; Hines et al., 2015; Siani et al., 2013; see Figure 17e). Furthermore, no evidence currently exists for B-Atm radiocarbon ventilation ages >3000 <sup>14</sup>Cyrs anywhere in the deep ocean after the B-A (see Figure 13), making it hard to account for the Baja California anomalies observed during the YD, even by appealing to an abyssal ocean source.

An alternative, and speculative, proposal might be that the radiocarbon anomalies observed off Baja California represent the influence of a radiocarbon-depleted sedimentary/geological carbon source, delivered to pore-waters during HS1 and the YD. Indeed, this hypothesis has been advanced to account for the far more “extreme” radiocarbon anomalies observed off Galapagos and the Gulf of California (Bova et al., 2018; Rafter et al., 2018, 2019; Stott et al., 2009). The proposal of a geological carbon source that supplemented the deep ocean's respired/disequilibrium dissolved inorganic carbon pool from as early as ~25 ka BP until approximately the end of the YD (Rafter et al., 2019; Ronge et al., 2016; Stott et al., 2019) is intriguing, and raises something of a carbon cycle puzzle. There are three key aspects to this conundrum: (a) the reconciliation of carbon isotope-, carbonate chemistry- and carbon cycle impacts across the last deglaciation; (b) the explanation of the different timing and patterns of various putative “smoking gun/volcano” records; and (c) the implications for long-term ocean-atmosphere carbon budget evolution across recurring glacial cycles.

The first of these items requires that the proposed addition of geological carbon, directly to the deep ocean, was sufficient to alter the deep ocean's radiocarbon signature, while leaving its δ<sup>13</sup>C signature unaffected (Rafter et al., 2019; Stott et al., 2019), and while also leaving sea floor carbonate sediment preservation unaffected (Marchitto et al., 2007; Rafter et al., 2019). The latter requirement stems from the fact that the addition of CO<sub>2</sub> to the deep ocean would be expected to lower pH and shift the dissolved carbonate system equilibria away from carbonate ion, causing carbonate sediments to dissolve wherever carbonate ion concentrations dropped below carbonate saturation levels (they are close to saturation throughout the modern deep Pacific [Key et al., 2004]).
<figure><img src="image_22.png" /><figcaption>**Figure 17.** Marine radiocarbon ventilation records that exhibit evidence for extreme radiocarbon depletion, potentially suggestive of for example, localized “geological” carbon sources to the sedimentary environment, or diagenetic overprinting. (a) Global average carbonate sediment mass accumulation rates (Cartapanis et al., 2018). (b) Atmospheric CO₂ (Monnin et al., 2001). (c) Atmospheric radiocarbon activity (black line, Intcal13 (Reimer et al., 2013); red line Intcal20 [Reimer et al., 2020]). (d) Baja California records (Lindsay et al., 2015, 2016; Marchitto et al., 2007), exhibiting two transient radiocarbon depletions, during HS1 and the YD. (e) Cubic spline fits (with shaded 95% confidence interval) for data from the deep North Pacific (>2 km, see caption for Figure 11), and the upper Southern Ocean (<2 km, see caption for Figure 13), compared with planktic and benthic records from Baja California. (f) Gulf of California records based on several different benthic foraminifer species and based on planktonic foraminifera (Rafter et al., 2018, 2019). (g) Galapagos foraminifer-based records (red filled circles, planktic foraminifera [Stott et al., 2009]; open circles and lines, benthic foraminifera [Bova et al., 2018; Stott et al., 2009]); compared with coral-based data from the Galapagos Plateau (orange open stars and line) that show no extreme depletions (Chen et al., 2020). Vertical shaded bars indicate the approximate timing of Heinrich Stadial 1 (HS1) and the Younger Dryas (YD). All ventilation ages have been “corrected” according to the difference between the Intcal20 and Intcal13 atmospheric radiocarbon curves.</figcaption></figure>
However, there is no evidence for significant carbonate dissolution at any of the eastern Pacific sites (Lindsay et al., 2015; Marchitto et al., 2007; Ortiz et al., 2004; Rafter et al., 2019; Stott et al., 2009). Globally, and across the Pacific, carbonate mass accumulation rates actually increased during the last deglaciation (Cartapanis et al., 2018; Figure 17a), as did carbonate ion concentrations in the deep Pacific (Allen et al., 2015, 2020; Marchitto et al., 2005; Yu et al., 2013). This would mean that significant and widespread geological CO₂ neutralization, either in the water column or at the sediment/sea-water interface, probably did not occur during deglaciation. Furthermore, if geological CO₂ neutralization by carbonate sediment dissolution did occur, it would cause ocean alkalinity to increase, and average ocean pCO₂ to decrease, thus counteracting deglacial CO₂ rise on a multi-millennial timescale.

To get around these “carbonate system” problems, a metamorphic source of bicarbonate ion has been hypothesized (Rafter et al., 2019). Similarly, if geological CO₂ was supplied in the ocean via deep sedimentary/metamorphic seeps in the south Pacific (Stott et al., 2019), with vertical fluid flow occurring over several hundred meters or more, one might also expect geological CO₂ to be neutralized by carbonate dissolution occurring deep in the sediment column, far below the sea floor. This workaround could potentially reconcile the radiocarbon observations from Galapagos (Bova et al., 2018; Stott et al., 2009) and the Gulf of California (Rafter et al., 2018, 2019) (and maybe the South Pacific [Ronge et al., 2016]) with records of deglacial carbonate chemistry and carbonate preservation (Allen et al., 2015, 2020; Cartapanis et al., 2018; Marchitto et al., 2005; Yu et al., 2013). However, it would also render them somewhat irrelevant to deglacial atmospheric CO₂ change (Rafter et al., 2019). This is because the supply of carbon in the form of bicarbonate would increase deep ocean DIC and ALK in a 1:1 ratio, thus decreasing the pCO₂ of the affected seawater only very slightly (Zeebe & Wolf-Gladrow, 2001). On a multi-millennial timescale, when this radiocarbon-dead bicarbonate-enriched water was transported to the sea surface by the overturning circulation, the “geological” contribution to reduced pH and increased pCO₂ in upwelling regions would therefore be negligible.

If they existed, submarine geological carbon sources (in the form of bicarbonate) could therefore be important for certain deglacial radiocarbon records, and perhaps for global carbon budgets over long time-spans via impacts on alkalinity, but not for deglacial atmospheric CO₂ change. The suggestion that a local volcanic input may be detected in sea floor sediments, and indeed seawater, without this input representing a major influence on the global radiocarbon distribution and budget, is hinted at by high resolution and high precision hydrographic measurements from the modern East Pacific rise (Jenkins et al., 2018). Such geological sources of carbon to the ocean interior are surely important for refining our understanding of the chemical inventories of the ocean; but they are too small to leave their mark on the global pattern of radiocarbon ventilation ages in the modern ocean (Broecker & Peng, 1982; Key et al., 2004). The latter are overwhelmingly controlled by air-sea gas exchange at the sea surface, and transports in the ocean interior (Koeve et al., 2015).

Accounting for the timing of the proposed volcanic contributions presents a further puzzle, since deglacial ice-volume and sea level changes (which were initiated ~19 ka BP [Lambeck & Chappell, 2001]) could not have triggered the submarine volcanic/metamorphic CO₂ release that are proposed to have occurred well before the LGM, and at least as early as 25 ka BP in some places (Ronge et al., 2016; Stott et al., 2019). Indeed, increased submarine volcanism would not be triggered by glacioeustatic sea level rise at all, but rather by glacioeustatic sea level fall (in opposition to subaerial volcanism) (Huybers & Langmuir, 2017). Any geological carbon release related to submarine volcanism during deglaciation would therefore need to reflect a rather auspiciously delayed response to sea level drop during the onset of glaciation before the LGM (Huybers & Langmuir, 2017), perhaps ~32 ka BP (Grant et al., 2012), or ~22 ka BP (Yokoyama et al., 2018). An alternative to the volcanic mechanism is the proposal of a CO₂-hydrate “capacitor,” possibly regulated by deep water temperature change (since increasing sea-level and pressure would on the contrary stabilize clathrates during deglaciation) (Stott et al., 2019). Although the details of this purely hypothetical mechanism have yet to be worked out, it would have the advantage of allowing geological carbon release to be linked more directly to local deglacial climatic/hydrographic triggers, rather than global glacioeustatic sea level.

The identification of extreme radiocarbon depletions at various locations and at various times between ~25 and 10 ka BP, might be taken to represent a “fly in the ointment” for the canonical account of deglacial marine radiocarbon change as primarily a story of evolving ocean circulation and air-sea gas exchange. Given the great efforts that have gone into assessing and replicating the evidence at various locations (e.g., Lindsay et al., 2015, 2016; Marchitto et al., 2007; Rafter et al., 2019), and in trying to place some of these results into an explicit geological
context (e.g., Stott et al., 2019), the evidence cannot be cast aside lightly. Indeed, a “geological” contribution to deglacial radiocarbon evolution in the marine sedimentary environment appears to be an incontrovertible conclusion at a few locations (Bova et al., 2018; Rafter et al., 2019; Stott et al., 2009, 2019). However, the further inference of a significant contribution from geological carbon sources to the ocean's dissolved inorganic carbon pool, seems more difficult to defend. This conclusion is supported in respect of the “global ocean” by the bulk of available data from the LGM (e.g., Skinner et al., 2017), and assembled here for the last deglaciation (Figures 13, 14, and 16). In respect of intermediate waters in the eastern low-latitude Pacific, it is further supported by U-series dated corals, from 627 m water depth on the Galapagos plateau (Chen et al., 2020), which show no evidence for extreme radiocarbon depletion in a region that is currently influenced by AAIW (stars in Figure 17f). As shown in Figure 15, these data exhibit a similar trend to coral data from the Southern Ocean (Burke & Robinson, 2012; Chen et al., 2015; Li et al., 2020), consistent with the expected preponderance of southern sourced mode waters in both of these locations, as in the modern ocean (Fiedler & Talley, 2006). The collected evidence therefore suggests that extreme radiocarbon depletion observed in benthic and planktonic foraminifera at several locations at the LGM and across the last deglaciation most likely reflects the influence of geological carbon sources on the sedimentary environment in which the foraminifera lived or were deposited, rather than the ocean interior per se.

Even if perceived as a “red herring,” as far as the main controls on deglacial carbon cycle evolution are concerned, the issue of geological carbon sources and their contribution to extreme radiocarbon depletion during deglaciation may yet serve as a useful tonic for our canonical account of Pleistocene carbon cycle dynamics. In this regard, it is interesting to note that glacial-interglacial variation of (subaerial) volcanic CO₂ flux, equivalent to ~30% of long-term average global emission rate, has emerged as an important “tuning parameter” (i.e., one that is adjusted to keep the surface carbon budgets balanced over time) in numerical model simulations of glacial-interglacial CO₂ that are forced only by insolation pacing of Earth system feedbacks (Ganopolski & Brovkin, 2017). The importance of sedimentation and sediment-seawater/pore-water interactions has also been underlined in modeling studies that implement sediment-modules and treat the joint ocean-atmosphere carbon reservoir as an “open” system (e.g., Tschumi et al., 2011). Something of a “geological turn” in emerging accounts of glacial-interglacial carbon cycle change is therefore apparent (Broecker et al., 2015; Ganopolski & Brovkin, 2017), and is arguably warranted, since the ocean-atmosphere system does not form a “closed” carbon budget on timescales longer than a few millennia (Jeltsch-Thömmes & Joos, 2020). Nevertheless, it seems clear that glacial-interglacial CO₂ and deglacial marine radiocarbon evolution is not either a simple story of solid Earth dynamics: the submarine (as opposed to sub-aerial) “geological” contribution to deglacial atmospheric CO₂ rise seems particularly challenging to explain when consideration is given to carbonate chemistry constraints, as well as the precise timing and structure of CO₂ changes. This includes their consistent link with the interhemispheric “bipolar seesaw” pattern, and associated changes in the AMOC, sea ice and ocean interior oxygenation for example (Gottschalk et al., 2016, 2019; Henry et al., 2016; Jaccard et al., 2016; Schüpbach et al., 2018). Identifying a “geological” contribution to global radiocarbon- and carbon cycle changes across glacial cycles, acting in parallel with the already well-defined (if also poorly quantified) contribution from the biological pump and air-sea gas exchange (Galbraith & Skinner, 2020), remains a further challenge for the future. Progress in addressing this challenge will benefit from a more detailed elucidation of the spatial and temporal evolution of seawater and sedimentary/authigenic-phase radiocarbon activities since the late glacial, in particular as a means of constraining isotope enabled numerical model scenarios that include carbonate chemistry and sedimentary fluxes.

## 6. Summary and Outlook

Radiocarbon is a particularly useful tool for the geosciences, applicable as a diagnostic carbon cycle tracer in paleoceanographic studies and in numerical models, and under certain circumstances as a radiometric dating tool extending to ~50 ka BP. A radiocarbon measurement performed on a single fossil entity can be used for one of three purposes (assuming that the fossil was in equilibrium with its parent reservoir at the time of formation):

1. to determine the radiocarbon activity of the fossil's parent reservoir at the time it formed, requiring knowledge of the true calendar age of the fossil entity;

2. to determine the time of formation of the fossil entity (i.e., its calendar age), requiring knowledge of the radiocarbon activity of its parent reservoir, and how it evolved over time;
3. to determine the extent of disequilibrium between the fossil entity's parent reservoir, at the time of the fossil entity's formation, relative to another contemporaneous reservoir (e.g., the ocean relative to the atmosphere at 19 ka BP, etc.).

The use of radiocarbon as a dating tool is severely limited by: past radiocarbon production changes; past changes in the exchange rates between the various Earth surface reservoirs; and variability within reservoirs, such as the ocean in particular. Uncertainty regarding past “marine reservoir ages” currently represents a major impediment to accurate calendar age dating of marine fossils using radiocarbon.

A further significant challenge that has been emphasized by the development of gas-source AMS capable of measuring very small carbon samples, is the elucidation of life-habit and taphonomic biases that may affect all fossil entities in the geological record. In this context, a pragmatic balance must be struck between cognizance of the distracting noise and bias that may plague radiocarbon data, versus the underlying signals that are informative, where “noise” and “signal” will always be defined relative to a particular question of interest.

Reliable (proxy) measurements of the radiocarbon activity of a parcel of water in the ocean interior will reflect the combined effects of variable radiocarbon production, the radiocarbon distribution at the sea surface and the transport trajectories and transit times for water moving through the ocean interior (Koeve et al., 2015), with minor contributions from newly remineralized (“old”) DOC and (“young”) organic carbon and biogenic carbonate (Walker et al., 2016). Marine radiocarbon activity therefore includes primarily the combined effects of air-sea gas exchange, ocean transport, and mixing; it does not provide a straightforward and immediate measure of water transit times.

Through its link to the ocean's disequilibrium- and respired carbon inventories (both relatively depleted in radiocarbon as compared to the atmosphere), marine radiocarbon maintains an intimate link with ocean-atmosphere carbon partitioning, and therefore the ocean's potential role in atmospheric CO₂ change. A larger sequestered carbon inventory in the ocean will typically imply a more radiocarbon depleted marine carbon reservoir, though this link will be modulated by ocean carbonate chemistry and marine biology.

Understanding and interpreting marine radiocarbon as a carbon cycle tracer still has much to gain from the implementation of this isotope in numerical models. Particularly promising areas of development are in the use of “offline” or “transport matrix” methods for efficiently resolving equilibrium solutions that would otherwise require very long integration times (Bardin et al., 2014; Khatiwala et al., 2005), as well as the development of “inverse methods” for application to radiocarbon in combination with other independent tracers (LeGrand & Wunsch, 1995). Due to the relatively short half-life of radiocarbon, its use in paleoceanography is restricted to the latter half of the last glacial cycle. In this context, most work has so far focused on the LGM and the last deglaciation. The analytically challenging, yet potentially rewarding, investigation of marine radiocarbon prior to the LGM represents an apparent gap to be filled. Similarly, detailed studies of marine radiocarbon evolution across the Holocene remain sparse, particularly for the deep ocean.

The great bulk of work looking at the distribution of radiocarbon in the past ocean has focused on the LGM (~20 ka BP), when atmospheric CO₂ was ~90 ppm lower than pre-industrial, and a very different climate state prevailed. The accumulated results demonstrate a significant “aging” of the ocean interior relative to the contemporary atmosphere (especially >2,000 m water depth, and by perhaps more than ~600 ¹⁴C yrs equivalent on average [Skinner et al., 2017]), thus supporting a role for ocean “ventilation” in CO₂ drawdown at the LGM, where ventilation is defined as including impacts on *both* disequilibrium- and respired carbon (i.e., gas exchange and ocean overturning rates [Galbraith & Skinner, 2020]). Although contributions to CO₂ drawdown at the LGM due to gas-exchange and ocean overturning are both supported by existing observations, determining their relative contributions remains a key challenge to address. The use of numerical models to this end will require that emphasis is placed on assessing changes in the “preformed” radiocarbon signatures of water in the ocean interior, in order to separate gas-exchange versus transport influences on radiocarbon and carbon sequestration (Eggleston & Galbraith, 2018). Surface reservoir age offsets, thus emerge as a particularly important target for reconstruction (Skinner, Muschitiello, & Scrivner, 2019), and for simulation (Butzin et al., 2017).

A long-standing overarching challenge remains to fully reconcile past atmospheric radiocarbon variability with reconstructed changes in radiocarbon production rates, and evolving marine radiocarbon activity (Dinauer et al., 2020). This challenge provides a stringent test, and an excellent training ground, for our understanding
of global carbon cycle evolution over the last glacial cycle, as it requires accurate and consistent descriptions of radionuclide production, ocean ventilation, terrestrial biomass evolution, and marine sedimentation. The eventual reconciliation of marine- and atmospheric radiocarbon change across the last deglaciation (since ~20 ka BP), has a direct bearing on the causes of natural atmospheric CO₂ variability, including 10–40 ppm centennial/millennial scale “jumps,” and the longer-term drift of CO₂ across recurrent glacial cycles that might arise from “open system” effects (Chalk et al., 2017; Hönisch et al., 2009).

Existing marine radiocarbon data suggest that the ocean's role in atmospheric CO₂ rise across the last deglaciation involved a pair of “ventilation seesaws” (between the North Atlantic and Southern Ocean [Broecker, 1998; Menviel et al., 2018; Skinner et al., 2014], and between the North Atlantic and North Pacific [Freeman et al., 2015; Max et al., 2014; Menviel et al., 2014; Okazaki et al., 2010]), as well as a pair of “mode changes” (in Southern Ocean convection from the onset of HS1 [Skinner et al., 2010], and in the AMOC from the onset of the B-A [Barker et al., 2010; Chen et al., 2015; Robinson et al., 2005; Skinner et al., 2013]). These “ventilation seesaws” and “step changes” were intimately connected to AMOC perturbations, through the thermal bipolar seesaw and its globally distributed impacts. The Pacific, and especially the Indian basin, stand out as major observational gaps to fill. Intermediate depths (900–2,000 m) in these basins may deserve special attention as potentially the primary locus for enhanced ocean-atmosphere equilibration and the exchange of carbon, and heat, during periods of “bipolar seesaw” activation (Baggenstos et al., 2019; Skinner et al., 2020).

A further major challenge that stands out in the context of deglacial radiocarbon evolution is an elucidation of the origin and carbon cycle implications of “extreme” radiocarbon depletions that have been observed between ~25 and ~10 ka BP at a number of sites. A dominant theory is that these represent radiocarbon-dead “geological” carbon injection at different times and at various locations on the sea floor (Rafter et al., 2019; Stott et al., 2019). However, the reconciliation of such a carbon source with marine carbonate preservation constraints seems to ensure that, even if it existed, it would have made little contribution to deglacial CO₂ rise. Furthermore, emerging coral radiocarbon data strongly question the validity of the foraminifer-based reconstructions as representative of ocean interior, as opposed to sedimentary/pore-water, phenomena (Chen et al., 2020). It therefore seems likely that the extreme radiocarbon depletions, at best, represent the impacts of “isotopic smoking guns” associated with small and localized geological carbon sources. Such radiocarbon-dead geological carbon sources can be observed, albeit with difficulty, in the modern ocean (Jenkins et al., 2018). Nevertheless, the apparent variability of such geological carbon sources, and their coordination with glacial-interglacial hydrographic and/or glacioeustatic change, underlines the importance of “open system” effects for the evolving alkalinity, carbon and radiocarbon budgets of the ocean-atmosphere system across recurrent glacial cycles.

In the 60 years since Willard Libby was awarded the Nobel Prize in chemistry “for his method to use carbon-14 for age determination in archeology, geology, geophysics, and other branches of science”, radiocarbon has proven to be far more useful than as just a dating tool. Far from having already borne its most significant intellectual fruits, radiocarbon is arguably a geochemical tracer that is poised to yield transformative progress throughout the geosciences. Annually resolved records of atmospheric radiocarbon are changing our understanding of solar dynamics (Miyake et al., 2013), and helping to improve even our best chronologies (Adolphi et al., 2018); novel analytical techniques are opening up new possibilities for high-resolution component/micro-analysis (Fagault et al., 2019; Wacker et al., 2013); a diversity of proposals regarding past radiocarbon production changes (Channell et al., 2018), as well as new reconstructions of past atmospheric radiocarbon (Cheng et al., 2018), have reset the stage for understanding the geodynamo and the biogeochemical influence of evolving sedimentary carbon fluxes; and ongoing careful data collection continues to shine new light on the multi-faceted role of the ocean as a regulator of atmospheric CO₂, emphasizing its physical transports as well as its climatically sensitive exchanges with both the atmosphere and the solid Earth. Much has yet to be learned about the Earth system through the use of this unique radio-isotope.

## Data Availability Statement

No new data were used in the review article, which is based on existing data from previously published sources. The sources of the published data used in the figures, and included in supplementary Table S1, are detailed in the relevant figure captions.
Acknowledgments

The authors are deeply indebted to many colleagues who have helped improve this review, through comments and discussion, and who have made available their data and model outputs for inclusion in figures. The authors are grateful for the inspiration provided, in different ways, by the late Nick Shackleton and the late Wally Broecker. Tom Marchitto, Piet Grootes, and two anonymous reviewers are thanked for their insightful and critical comments, which helped to improve an earlier draft of this review. The authors are indebted to Eric Galbraith and Francois Primeau for input on numerical modeling of radiocarbon, as well as Laurie Menviel, Peter Köhler, Mathis Hain, Ayako Abe-Ouchi, Megumi Chikamoto, Fortunat Joos, Andreas Schmittner, Juan Muglia, Samar Khatiwala, Ashley Dinauer, Andy Ridgwell, Guillaume Soulet, Maria de la Fuente, and a great number of other paleoceanographer colleagues for the provision of model outputs, data, and discussion. This review benefited from stimulating discussions during INQUA IPODS, PAGES OC3, and Intcal marine working group meetings. L. C. Skinner acknowledges funding from the Royal Society, the Isaac Newton Trust, and NERC grant NE/L006240/1, which supported data collection that has informed this review. E. Bard is supported at CEREGE by EQUIPEX ASTER-CE-REGE and ANR MARCARA.

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