The Paleogene history of biogenic opal accumulation in the North
Atlantic provides insight into both the evolution of deepwater circulation
in the Atlantic basin and weathering responses to major climate shifts.
However, existing records are compromised by low temporal resolution and/or
stratigraphic discontinuities. In order to address this problem, we present
a multi-site, high-resolution record of biogenic silica (bioSiO2) accumulation from Blake Nose (ODP Leg 171B, western North Atlantic) spanning the early Paleocene to late Eocene time interval (∼65–34 Ma). This record represents the longest single-locality history of marine bioSiO2 burial compiled to date and offers a unique perspective into changes in bioSiO2 fluxes through the early to middle Paleogene extreme greenhouse interval and the subsequent period of long-term cooling. Blake Nose bioSiO2 fluxes display prominent fluctuations that we attribute to variations in sub-thermocline nutrient supply via cyclonic eddies associated with the Gulf Stream. Following elevated and pulsed bioSiO2 accumulation through the Paleocene to early Eocene greenhouse interval, a prolonged interval of markedly elevated bioSiO2 flux in the middle Eocene between ∼46 and
42 Ma is proposed to reflect nutrient enrichment at Blake Nose due to
invigorated overturning circulation following an early onset of Northern
Component Water export from the Norwegian–Greenland Sea at ∼49 Ma. Reduced bioSiO2 flux in the North Atlantic, in combination
with increased bioSiO2 flux documented in existing records from
the equatorial Pacific between ∼42 and 38 Ma, is interpreted
to indicate diminished nutrient supply and reduced biosiliceous productivity at Blake Nose in response to weakening of the overturning circulation. Subsequently, in the late Eocene, a deepwater circulation regime favoring limited bioSiO2 burial in the Atlantic and enhanced bioSiO2 burial in the Pacific was established after ∼38 Ma, likely in conjunction with re-invigoration of deepwater export from the North Atlantic. We also observe that Blake Nose bioSiO2 fluxes through the middle Eocene cooling interval (∼48 to 34 Ma) are similar to or higher than background fluxes throughout the late Paleocene–early Eocene interval (∼65 to 48 Ma) of intense greenhouse warmth. This observation is consistent with a temporally variable rather than constant silicate weathering feedback strength model for the Paleogene, which would instead predict that marine bioSiO2 burial should peak during periods of extreme warming.
Introduction
Biogenic silica (bioSiO2) secretion by marine plankton and the
subsequent accumulation of biosiliceous marine sediments represent the main
output flux in the global silicon cycle (Tréguer and De La Rocha, 2013).
The present-day silicon cycle is also closely linked to the carbon cycle
because diatoms – the most successful and efficient
bioSiO2-secreting plankton group in the modern oceans – are also
the key marine primary producers, responsible for up to 40 % of total
global net photosynthesis per year (Smetacek, 1999). Owing to the ballast
effect of their siliceous valves, diatoms are extremely efficient at
exporting organic carbon (Corg) from the surface to the deep ocean and facilitating Corg burial in marine sediments (Yool and Tyrrell, 2003). This relationship between the silicon and carbon cycles has profound
implications for understanding climate change on both long and short
timescales in the past, founded on the premise that sedimentary
bioSiO2 mass accumulation rates (fluxes) represent the rate of
bioSiO2 burial and, importantly, that the burial rate reflects
bioSiO2 production in surface waters at the time of deposition
(Ragueneau et al., 2000; Yool and Tyrrell, 2005).
One fundamental control on marine siliceous plankton production is the
amount of dissolved silicon supplied to the oceans from terrestrial silicate
weathering, a chemical process that ultimately consumes atmospheric CO2 and releases silicic acid and alkalinity to the oceans (i.e., when combined with carbonate burial in the oceans; Walker et al., 1981; Fontorbe et al., 2020; Penman et al., 2020). By moderating atmospheric CO2, the silicate weathering feedback is postulated to operate as a thermostat, maintaining the Earth's surface within a habitable range of temperatures since early in geological history (Kasting, 2019). In today's rapidly warming world, an accurate understanding of the operation of silicate weathering as a climate feedback mechanism is essential.
Past transient greenhouse warming events, such as the “hyperthermal” events
of the early to middle Paleogene (Paleocene and Eocene epochs; ∼66–34 Ma), offer ancient points of comparison for the present-day warming
and future climate scenarios. Existing studies suggest that Paleogene
bioSiO2 accumulation patterns are directly linked to variations
in continental weathering on both long (Muttoni and Kent, 2007; Cermeño
et al., 2015; Renaudie, 2016) and short (Witkowski et al., 2014; Penman,
2016; Penman et al., 2019) timescales. Additionally, large volumes of
diatomite and diatom-rich clays deposited on continental shelves during the
early Paleogene (e.g., Oreshkina and Aleksandrova, 2007) suggest that the
supply of dissolved silicon from continental weathering under greenhouse
climates exerted a strong influence on marine bioSiO2 accumulation. These interpretations are based on an assumption that
diatoms were already key players in global silicon and carbon cycling in the
early Cenozoic (see Fontorbe et al., 2016; Conley et al., 2017). Testing
this assumption using the diatom fossil record is problematic due to the
vulnerability of diatom bioSiO2 to diagenetic alteration (see
Witkowski et al., 2020b, for a discussion and references). Thus, most
interpretations concerning the long-term silicon availability impact on
marine diatom production, as well as most scenarios for the timing of diatom
rise to ecological prominence, are based on a range of indirect evidence,
including modeling, isotope (bio)geochemistry, statistical treatment of
large databases, and insights from other biosilicifying groups (for a recent
synthesis, see Hendry et al., 2018).
In the modern oceans, deepwater circulation also exerts a major control on
marine bioSiO2 accumulation patterns throughout the ocean
basins. Firstly, ocean circulation determines the distribution and
concentration of limiting macronutrients (N, P, dissolved Si) in deep waters
and their upwelling into surface waters where they fuel primary production,
and, secondly, circulation impacts bioSiO2 preservation in
seafloor sediments (Ragueneau et al., 2000). Most of the present-day
biosiliceous production is focused along continental margins in areas where
diatoms can take advantage of nutrients supplied from continental runoff and
coastal upwelling (Malviya et al., 2016). A large proportion of
bioSiO2, however, is recycled even before settling out of the
photic zone (Van Cappellen et al., 2002), since the modern oceans are
undersaturated with respect to SiO2 at all depths. Only a fraction of
bioSiO2 produced in the photic zone therefore reaches the ocean
floor, and, furthermore, only a fraction of the exported bioSiO2
is incorporated into sediments and preserved (Frings, 2017).
Both the strength of the silicate weathering feedback and ocean circulation
patterns are believed to have undergone profound changes through the
early to middle Paleogene as the Earth transitioned from a hothouse, ice-free
climate state (e.g., Zachos et al., 2008; Kirtland-Turner et al., 2014;
Anagnostou et al., 2016) to an icehouse climate marked by continental-scale
ice sheets (Zachos et al., 2001; Miller et al., 2020). Isotopic weathering
proxies (87Sr/86Sr and δ7Li) display a broad minimum spanning the late Paleocene through early Eocene interval (Misra and
Froelich, 2012), which have been interpreted to indicate either flat
continental relief through this period (and thus reduced rates of
continental runoff; Froelich and Misra, 2015) or evidence for a variable
strength of the negative feedback between climate and silicate weathering
(Caves et al., 2016). Sea level fall associated with the onset of the
Antarctic glaciation at the Eocene–Oligocene transition (EOT) (Zachos et
al., 1996) and the intensification of the Himalayan orogeny are also
thought to have altered the dominant weathering regime by facilitating
physical rather than chemical weathering and by exposing large volumes of
fresh rock to erosion and weathering (Cermeño et al., 2015).
The early Cenozoic was a time of low thermal gradients between surface and
deep waters and between high and low latitudes, which limited vigorous
overturning circulation (Moore et al., 2008; Vahlenkamp et al., 2018). There
is, however, little consensus on the timing of the onset of production of Northern
Component Water (NCW) sourced from the high-latitude North
Atlantic, a precursor to today's North Atlantic Deep Water. Estimates
for NCW onset vary from the late early Eocene (∼49–50 Ma; Hohbein
et al., 2012) through the late Eocene (∼38 Ma; e.g., Borrelli
et al., 2014; Coxall et al., 2018) to across the EOT interval
(∼34 Ma; Via and Thomas, 2006; Abelson and Erez, 2017).
Regardless, the long-term global cooling spanning the middle and late Eocene
(∼48–34 Ma) is thought to have resulted in enhanced upwelling, and the opening of the Drake Passage is viewed as a milestone in establishing the global pattern of thermohaline circulation in its present-day form in the Atlantic (Via and Thomas, 2006; Katz et al., 2011; Borrelli et al., 2014; Abelson and Erez, 2017).
Despite the importance of siliceous biota in the present-day carbon cycle,
our understanding of the temporal trends in marine bioSiO2 accumulation through the early Paleogene is limited. First-order
observations indicate an association between peak chert–porcelanite
occurrence and deepwater temperatures through the Early Eocene Climatic
Optimum (EECO) (Muttoni and Kent, 2007; Witkowski et al., 2020b) and during
short-lived hyperthermal events of the early Eocene (Penman et al., 2019).
The rapid cooling at the end of the Eocene is also widely regarded as the
period of diatom proliferation and diversification, especially in the
Southern Ocean (Egan et al., 2013; Lazarus et al., 2014; Renaudie, 2016).
However, trends in marine bioSiO2 accumulation in the period
between these temporally broadly isolated events representing contrasting
climate states are not well documented. The longest currently available
perspective on marine bioSiO2 accumulation (Cretaceous through
Miocene) is based on Deep Sea Drilling Project (DSDP) Leg 1 through 44
smear-slide data (i.e., data gathered between 1968 and 1978) converted to
mass accumulation rates and binned into 10 Myr increments
(Miskell et al., 1985). Direct sediment measurements of bioSiO2 concentrations (with calculated fluxes) through the Paleogene are sparse
(e.g., Diester-Haass, 1995; Salamy and Zachos, 1999; Diekmann et al., 2004;
Lyle et al., 2005; Iwasaki et al., 2014) and mostly focus on restricted
time windows of the late Eocene through early Oligocene interval. A major
reason for this is that bioSiO2 is highly vulnerable to
water-column and seafloor dissolution, which results in early Paleogene
siliceous phytoplankton occurrences often being confined to narrow
stratigraphic intervals at many sites (see Barron et al., 2015; Witkowski et
al., 2020b).
The observations summarized above provoke three fundamental questions. (1) How did bioSiO2 burial flux evolve through the early to middle
Paleogene? (2) What were the main controls on changes in marine
bioSiO2 burial in this time interval? (3) What was the
bioSiO2 burial response to long-term Paleogene climate changes?
Through the early Paleogene, bioSiO2 accumulation was largely
focused in the Atlantic and on the Eurasian Platform (Miskell et al., 1985;
Muttoni and Kent, 2007; Moore et al., 2008; Barron et al., 2015; Wade et
al., 2020). Unusually expanded lower Paleocene through upper Eocene
biosiliceous successions were recovered from Blake Nose in the midlatitude
western North Atlantic (Shipboard Scientific Party, 1998a–f; Witkowski et
al., 2020a). In order to gain a quantitative insight into how
bioSiO2 burial evolved through the early Paleogene hothouse and
the ensuing period of global cooling, we have generated a composite
high-resolution Blake Nose bioSiO2 flux record from
∼65 to 34 Ma, spanning nearly the entire Paleocene and Eocene
epochs. This work follows on from two previous publications with a focus on
the Blake Nose early to middle Paleogene siliceous microfossils: (1) Witkowski
et al. (2020a), in which a revised chronological framework is proposed for
Sites 1050 and 1051; and (2) Witkowski et al. (2020b), in which Paleogene
trends in chert and porcelanite occurrences are compared to spatial and
temporal patterns in Atlantic biosiliceous sediment occurrences. In our
study here, we aim to determine the main controls on bioSiO2 fluxes in a key locus of biosiliceous accumulation in the western North
Atlantic (Blake Nose) during the early to middle Paleogene – a period of
Cenozoic climate change characterized by profound variations in global
temperature.
Materials and methodsStudy sites and stratigraphy
This study is focused on drill cores recovered as part of the Ocean Drilling
Program (ODP) Blake Nose Paleoceanographic Transect and includes Holes
1049A, 1050A/C, 1051A, 1052B/F, and 1053A (Shipboard Scientific Party,
1998b–f) (Fig. 1) (Table 1). The transect was drilled on Blake Nose (BN;
also often referred to as “Blake Ridge”) in the western North Atlantic Ocean in order to reconstruct the Cretaceous–Paleogene paleoceanographic history of the region adjacent to the South Atlantic Bight (Shipboard Scientific Party, 1998a), offshore of the southeastern US seaboard between Florida Straits and Cape Hatteras (Gula et al., 2016). The BN is a
northeast-trending extension of the Blake Plateau comprised of a Cretaceous
to Paleogene continental margin succession (Pinet et al., 1981; Shipboard
Scientific Party, 1998a) (Fig. 1b). As Paleogene sediments draping BN were
deposited on the seaward slope of a large reef formed in the Early
Cretaceous, the transect sites likely retain the relative depths of the
paleo-reef system (Shipboard Scientific Party, 1998a). Drilling at BN, as
well as at sites further north along the North American margin, documented
well-preserved early Paleogene siliceous microfossils (e.g., Gombos, 1982;
Nishimura, 1992; Hollis, 2014). The BN transect sites that recovered the
most expanded early to middle Paleogene sections (Sites 1050 and 1051) include
only a few narrow chert and porcelanite-bearing intervals (Witkowski et
al., 2020b) and sparse clinoptilolite (a zeolite alteration product of
biogenic silica) occurrences, which suggests minimal diagenesis of
sedimentary bioSiO2. As such, the good overall preservation of
siliceous microplankton in Paleogene sediments at BN, combined with the
exceptionally long stratigraphic span of the record recovered during Leg
171B (Witkowski et al., 2020a, and references therein), makes the BN
transect especially well-suited for reconstructing variations in
bioSiO2 burial flux in the early Paleocene through late Eocene
time period. To this end, we examined 1230 samples from five BN drill sites:
Sites 1049, 1050, 1051, 1052, and 1053.
Maps showing the location of sites considered in the present paper: (a) Blake Nose (red star) in the western North Atlantic and eastern
equatorial Pacific sites (black stars). Base map generated using Ocean
Drilling Stratigraphic Network (2021) Advanced Plate Tectonic Reconstruction
service (https://www.odsn.de/, last access: 12 April 2021). (b) Location of the Ocean Drilling Program (ODP) Leg 171B sites on Blake Nose. Modified from Shipboard
Scientific Party (1998a).
Sites included in this study, along with geographic coordinates,
site chapters, and number of samples examined for bioSiO2. Sites
used for bioSiO2 flux calculations are in italics.
ODP holeLatitudeLongitudeWater depthReferenceNumber of samples(m)examined forbioSiO21049A30∘08.5436′ N76∘06.7312′ W2656.1Shipboard Scientific Party (1998b)701050A30∘05.9977′ N76∘14.1011′ W2299.8Shipboard Scientific Party (1998c)2731050C30∘05.9953′ N76∘14.0997′ W2296.5Shipboard Scientific Party (1998c)71051A30∘03.1740′ N76∘21.4580′ W1982.7Shipboard Scientific Party (1998d)7621052B29∘57.0791′ N76∘37.6098′ W1345.0Shipboard Scientific Party (1998e)131052F29∘57.0794′ N76∘37.6094′ W1343.5Shipboard Scientific Party (1998e)261053A29∘59.5385′ N76∘31.4135′ W1629.5Shipboard Scientific Party (1998f)79Total samples examined 1230
Site 1049 is the most distal and deepest site included in this study
(1000–2000 m paleodepth; Shipboard Scientific Party, 1998b) (Fig. 1b; Table 1). The Paleogene section of Hole 1049A was poorly recovered due to the
extensive presence of chert horizons. As a consequence, numerous
biostratigraphic datums are poorly constrained through the recovered
sequence, and age control is only approximate, especially through the
early–middle Eocene transition (EMET). Based on the recent revisions to the
bio-magnetostratigraphy of Holes 1050A and 1051A, however, Witkowski et al. (2020b) proposed a revised age model for the Paleocene through Eocene
interval of Hole 1049A. Here, we examine 70 samples from the chert-free,
siliceous-microfossil-bearing interval of Hole 1049A, spanning Cores
1049A-3H through -12X (∼21.1 to 88.1 compacted meters below
seafloor; m b.s.f. – see Witkowski et al., 2020a; Table 1). Due to incomplete
recovery and the temporal patchiness of the record, however, we do not
include data from Hole 1049A in bioSiO2 flux calculations.
Site 1050 (1000–2000 m paleodepth; Shipboard Scientific Party, 1998c) was
drilled several kilometers upslope of Site 1049 (Fig. 1b; Table 1). The Paleogene
succession cored in Holes 1050A and 1050C is considerably more expanded and
stratigraphically more complete than in Hole 1049A. Siliceous microfossils
occur throughout the succession cored in Hole 1050A, but in Hole 1050C
siliceous microfossils are confined to Core 1050C-2R (Witkowski et al., 2020b). We apply the age model of Witkowski et al. (2020a), who interpreted
the presence of two major stratigraphic gaps. For this study, we examined
273 samples from Hole 1050A (Cores 1050A-2H through -36X; ∼11
to 319.3 compacted m b.s.f.) and 7 samples from Hole 1050C (Core 1050C-2R; 328 to 336 compacted m b.s.f.).
Site 1051, the intermediate-depth site of the BN transect (1000–2000 m
paleodepth; Shipboard Scientific Party, 1998d) (Fig. 1b; Table 1), recovered
the most expanded lower Paleocene through upper Eocene succession among Leg
171B sites. Siliceous microfossils occur throughout this succession, except
for within several narrow dissolution intervals (for details see Witkowski
et al., 2020b). We use the age model of Witkowski et al. (2020a), who showed
that the Hole 1051A succession is interrupted by two major gaps that are
broadly correlative with the hiatuses in Hole 1050A (see also Röhl et al., 2003). A total of 762 samples from the entire succession cored at Hole 1051A (∼8.5 to 644 compacted m b.s.f.) were examined for this study.
Site 1052 is the shallowest site of the BN transect sites, drilled near the
crest of the BN (Fig. 1b; Table 1). Most of the middle bathyal (600–1000 m
paleodepth; Shipboard Scientific Party, 1998e) Paleogene succession at this
site is truncated by a prominent hiatus. In this study, we include a narrow
composite interval of Holes 1052B and 1052F (∼77 to 131 meters composite depth, mcd) spanning the Middle–Late Eocene Turnover (MLET) (Kamikuri and Wade, 2012). As this interval overlaps parts of
Holes 1051A and 1053A, it is not considered in sediment flux calculations.
For age control at Site 1052, we use the bio-magnetostratigraphic
constraints from Shipboard Scientific Party (1998e), Ogg and Bardot (2001),
and Wade et al. (2012), following Witkowski et al. (2020b). A total of 39
samples from Site 1052 were used in this study.
Site 1053 was drilled between Sites 1051 (intermediate depth) and 1052
(shallowest depth) in the upper part of the BN transect (500–700 m
paleodepth; Shipboard Scientific Party, 1998f) (Fig. 1b; Table 1). Site 1053
recovered an expanded siliceous-microfossil-rich upper Eocene section with
no detectable stratigraphic gaps, as indicated by the age model of Borrelli
et al. (2014). At total of 79 samples from Hole 1053A (∼0.5
to 183 meters below seafloor, m b.s.f.) were examined for this study.
Despite two major discontinuities and multiple recovery gaps, our composite
BN record is comprised of data from five sites and spans the earliest
Paleocene (∼64.74 Ma; magnetochron C28n in Hole 1051A, Witkowski et al., 2020a) through latest Eocene (∼33.94 Ma; magnetochron C13r in Hole 1053A; Borrelli et al., 2014) interval. This composite represents the longest currently available single-locality record of deep-sea biosiliceous sedimentation through the Paleogene. We report all ages relative to the Gradstein et al. (2012) timescale, hereafter referred to as GTS2012.
bioSiO2 measurements
bioSiO2 concentrations were determined by means of a Hach DR-3900
spectrophotometer using Hach method 8186 (heteropoly blue method). All
spectrophotometric analyses closely followed the wet alkaline extraction
procedure of Olivarez Lyle and Lyle (2002). Unlike Olivarez Lyle and Lyle
(2002), however, for base extraction we used 1 M KOH and 10 mg ground
sediment subsamples rather than 2 M KOH and 20 mg subsamples. This was done
in order to avoid SiO2 polymerization (Annette Olivarez Lyle, personal communication, 2015), manifested by the precipitation of whitish filaments in test tubes following base extraction conducted at higher KOH
concentrations with larger subsamples. bioSiO2 concentrations for
individual sites are tabulated in Tables S1 through S5 in the Supplement, and
data used for bioSiO2 flux calculations are presented in Figs. S1–S3 in the Supplement.
Three methods were employed to monitor analytical precision of the
bioSiO2 measurements: (1) one sample in each analyzed batch was
subject to stepwise standard addition against a target curve using liquid
SiO2 standard supplied by Hach (average target curve R2=0.994, n=84); (2) one random sample from every sample batch was also analyzed in duplicate, with good correlation between duplicate analyses (average R2=0.98, n=92; Fig. S4); and (3) one of three in-house consistency standards was analyzed in approximately every
second sample batch.
Sediment mass accumulation rate calculations
All sediment mass accumulation rate (hereafter: flux) values in this work
are expressed as grams per square centimeter (g cm-2) per 1000 years (kyr) and are calculated using standard terms from previous studies (e.g., Diester-Haass, 1995; Piela et al., 2012; D'haenens et al., 2014).
MAR(gcm-2kyr-1)=sedimentarycomponent[gcomponent/gbulksediment]×linearsedimentationrate(LSR)[cmkyr-1]×sedimentdrybulkdensity(DBD)[gcm-3]
Use of magneto-biostratigraphic age models to establish LSRs typically
produces unrealistic jumps in sediment flux estimates at
magnetostratigraphic boundaries, with order-of-magnitude differences between
consecutive age–model tie points. In order to smooth out such abrupt
features, which we deem to be artifacts of the applied age models, in our
flux records, we fitted polynomial regressions against the age vs. depth
curves (or segments thereof comprised between hiatuses), following the
approach of Piela et al. (2012). The datasets developed in the present work
are based on several holes that include several hiatuses, which is why
robust age models that are consistent between holes are essential to obtain
a reliable composite stratigraphy. We therefore plot the flux records
derived from smoothed LSR estimates using ages interpolated from the
original (i.e., non-smoothed) age–depth curves (Figs. S1–S3).
Sediment flux studies often estimate wet bulk sediment density through
calibration of high-resolution estimates of wet bulk density (obtained via
gamma ray attenuation, GRA, analysis) against discrete DBD measurements
collected during routine shipboard analysis. In the present work,
establishing a single GRA–DBD correlation over the entire cored interval
proved ineffective for Sites 1050 and 1051, likely due to the downhole
increase in compaction. Instead, we estimated DBD for a given depth by
interpolating between shipboard discrete DBD measurements (Shipboard
Scientific Party, 1998b–d, f). Sediment density plots, LSRs, and calculated
fluxes are included in the Supplement (Figs. S1–S3 and Tables S2–S3, S5).
Stable isotope and pCO2 data
Our interpretation of possible controls on early to middle Paleogene
bioSiO2 accumulation is based on comparison to published isotopic
weathering (87Sr/86Sr, 187Os/188Os, and δ7Li; Ravizza et al., 2001; Ravizza and Peucker-Ehrenbrink, 2003; Misra
and Froelich, 2012; Klemm et al., 2005) and paleocirculation proxies
(δ13C, δ18O; Cramer et al., 2009), as well as a recent atmospheric pCO2 reconstruction (Foster et al., 2017) and silicate weathering flux model (hereafter SWF) (Caves et al., 2016). For further discussion and a full documentation of data sources see the Supplement (“Stable isotope and pCO2 data” section and Fig. S5).
Statistical treatment
Smoothed long-term trends in bioSiO2 flux and published geochemical records were obtained via local regression (abbreviated as LOESS; Cleveland et al., 1992) computed using R Studio v. 3.5.1. Statistical analysis was performed on smoothed time series (bioSiO2 flux, δ13C, δ18O, pCO287Sr/86Sr, 187Os/188Os, δ7Li, and SWF) using Statistica 13.1 package.
The degree of covariance of the analyzed variables was assessed by
correlation analysis. A normality test procedure was carried out for all
variables using the Shapiro–Wilk test (α=0.05). The Pearson
correlation coefficient was used to assess covariance for each pair of
variables characterized by a normal distribution. The nonparametric
Spearman correlation coefficient was used when non-normal distribution was
indicated for a given variable by the Shapiro–Wilk test.
The analysis also involved a multiple regression model, which describes the
relationship of the dependent variable Y with a set of independent variables X1, X2,…,Xk (which, in this study, is the
relationship between bioSiO2 flux and other proxy records). It is
defined by Eq. (1):
Y=β0+β1X1+β2X2+…+βkXk+ξ,
where βj represents model parameters (regression coefficient), and
ξ is a random component. The parameters of the regression equation are estimated using the least-squares
method, and the determination coefficient and standard error of estimation are
used to assess the goodness of the model.
Results and interpretation
Our new composite %bioSiO2 record from Blake Nose spans the
interval between ∼65 and 34 Ma (Fig. 2a), representing the
longest single-locality record of bioSiO2 concentrations compiled to date. The composite record, however, lacks data in two short time
windows: between ∼53.5 and 52.0 Ma (magnetochrons C24n through C23n) and between ∼47.5 and 49.0 Ma (i.e., through the EMET). This is due to the presence of prominent hiatuses at all study
sites spanning these intervals (Shipboard Scientific Party, 1998b–d;
Witkowski et al., 2020a) (Fig. 2a).
Composite Blake Nose weight percent biogenic silica concentrations (a) and fluxes (b) through the early to middle Paleogene plotted against key tectonic and climatic events (c), silicate weathering flux as modeled by Caves et al. (2016) (d), global benthic foraminiferal δ18O compilation of Cramer et al. (2009; rescaled to GTS2012) (e), and
pCO2 reconstruction (Foster et al., 2017) and 87Sr/86Sr ratios (Misra and Froelich, 2012) (f). The schematic representation of climatic trends next to the chronostratigraphy panel is consistent with Cramwinckel et al. (2018). Abbreviations: wt % – weight percent; bioSiO2 – biogenic silica; GTS2012 – Geologic Time Scale 2012; see Gradstein et al. (2012).
The BN %bioSiO2 composite shows variable but generally high
values between ∼65 and 49 Ma (Fig. 2a). Two broad %bioSiO2 maxima are observed within this high-bioSiO2 interval, culminating at ∼61.5 Ma and at ∼51.5 Ma (Fig. 2a). These maxima are separated by a broad low in
%bioSiO2 with a nadir centered at approximately the
Paleocene–Eocene boundary (∼56 Ma). From ∼49 Ma to the end of the record at ∼34 Ma, %bioSiO2 levels are considerably lower and less variable (Fig. 2a), with a distinct maximum culminating at ∼44 Ma.
Long-term trends in BN bioSiO2 fluxes are calculated based on a
composite record built from datasets generated from Sites 1050, 1051, and
1053. Through the Paleocene and early Eocene, Site 1050 generally displays
lower bioSiO2 fluxes than Site 1051 (Fig. 2b). From
∼46 to 34 Ma, both bioSiO2 flux trends and values
are remarkably consistent between Sites 1050 and 1051 (Fig. 2b). The short
time interval in which records from Site 1051 and Site 1053 overlap also
reveals coherent bioSiO2 flux values (Fig. 2b). Thus, following a
period of high inter-site variability through the Paleocene, three intervals
of elevated bioSiO2 fluxes are observed, which are consistent
between sites and peak at ∼53.2, ∼43.3, and ∼37.7 Ma. The overall patterns in %bioSiO2 and bioSiO2 flux estimates are also consistent, especially through the middle and late Eocene. Most importantly, however, the bioSiO2 flux values fall within the same order of magnitude through most of the study period (except for peak
bioSiO2 fluxes at Site 1051 between 54 and 53 Ma and from 44 to 43 Ma). Furthermore, our record consistently shows that bioSiO2
fluxes through the middle Eocene cooling were, on average, higher than (Site
1050) or similar to (Site 1051) bioSiO2 fluxes through the early
Eocene period of extreme greenhouse warmth (Fig. 2b).
Impact of hiatuses and diagenesis on BN bioSiO2 flux estimates
The bioSiO2 records from Sites 1050 and 1051, which constitute
the older part of the composite presented here, are interrupted by hiatuses.
These discontinuities in the BN record could introduce a bias to the flux
estimates, for instance by influencing the LSR calculations. The age models
for Holes 1050A/C and 1051A used in this study (see Witkowski et al., 2020a,
for details), however, are highly consistent in that the hiatuses are
identified in correlative intervals, and, furthermore, LSRs used in flux
calculations were subjected to polynomial smoothing, which should eliminate
most short-term artifacts imposed by age model imperfections.
bioSiO2 flux estimates could also be compromised by winnowing,
which could concentrate biosiliceous particles over some areas of the
seabed, while removing them from adjacent areas. In the core description
logs for BN sites included in the present bioSiO2 flux
reconstruction, explicit mention of winnowing is made only in one instance,
i.e., for Core 1051A-41X (Shipboard Scientific Party, 1998d). This core
is also characterized by the abundant presence of zeolite crystals (likely
clinoptilolite; Jakub Witkowski, unpublished observations), which are an indicator of bioSiO2 diagenesis (see Fenner, 1991). For this reason, Core 1051A-41X was excluded from the present study. Also, %bioSiO2 measurements were not performed on the sparse cherty or porcellanic intervals at Sites 1050 and 1051. Scanning electron microscope examination of diatoms from the remaining intervals of the BN composite indicates only minor diagenetic effects on the siliceous microfossils, manifested mostly by dissolution of the most delicate parts of the valves, such as areole occlusions or pore fields. For these reasons, the bioSiO2 fluxes reconstructed in this study are deemed to be robust.
Controls on bioSiO2 accumulation through the early to middle Paleogene at Blake Nose
bioSiO2 production, export, and preservation in marine sediments
are influenced globally by dissolved silicon supply to the oceans derived
from terrestrial weathering, which is closely linked to climate via a
negative feedback (e.g., Walker et al., 1981), and by ocean circulation
patterns and upwelling, which supply the bulk of macronutrients to surface
waters (Miskell et al., 1985; Handoh et al., 2003). In order to gain insight
into the influence that each of these factors has exerted on
bioSiO2 accumulation through the early to middle Paleogene at BN, we
compared the bioSiO2 flux composite record to published composite
global benthic foraminiferal δ18O and δ13C
records, pCO2 proxy estimates, proxy records of continental weathering (87Sr/86Sr, 187Os/188Os and δ7Li), and modeled silicate weathering flux (SWF) (Figs. 2, S5).
We find that bioSiO2 flux is moderately correlated with modeled SWF
(r=0.597, p<0.05) and more strongly, but inversely, correlated with
both pCO2 (r=-0.775, p<0.05) and δ18O
(r=-0.618, p<0.05) (Fig. 3). A weaker, but still statistically
significant, correlation exists between bioSiO2 flux and δ7Li (r=-0.473, <0.05), 187Os/188Os
(r=0.430, p<0.05), and 87Sr/86Sr (0.418, p<0.05) (Fig. 3). No statistically significant correlation has been found
between bioSiO2 flux and trends in benthic foraminiferal δ13C. These relationships suggest that, overall, through the
early to middle Paleogene, bioSiO2 flux at BN was indirectly shaped
by a combination of changes in atmospheric greenhouse gas levels, bottom
water temperatures (assuming ice-free poles through our study period), and
supply of solutes from terrestrial silicate weathering.
Correlation scatter plots for the three strongest statistical relationships identified in the present study: biogenic silica flux vs.
δ18O(a), pCO2(b), and silicate weathering flux (c). Abbreviations: bioSiO2 – biogenic silica; SWF – silicate weathering flux. Data sources indicated in the text.
Multiple regression indicates four significant variables shaping BNbioSiO2 flux: δ18O, pCO2, δ13C, and 87Sr/86Sr. Except for δ13C, this is consistent with the correlations discussed above. Notably, SWF was excluded by the multiple
regression model. This is likely due to the high overall similarity in
temporal trends displayed by BN bioSiO2 flux and SWF. The
multiple regression model equation takes the form
bioSiO2flux=798.57-0.156×δ18O-0.0008×pCO2-0.111×δ13C-1126.58×87Sr/86Sr.
This model explains ∼71 % of BN bioSiO2 flux
variance, with a standard error of estimation equal to 0.09. We find that
the model reproduces our calculated bioSiO2 flux values reasonably
well (Fig. 2b), suggesting that the use of smoothed datasets is suitable for
identifying long-term trends in bioSiO2 fluxes. Thus, both the
correlations and multiple regression suggest that BN bioSiO2 flux
was shaped mostly by δ18O, pCO2, and the supply of continental weathering products – all of which are related to the
temperature–silicate weathering feedback.
DiscussionImplications for paleocirculation
The Blake Nose area is positioned on the western margin of the North
Atlantic subtropical gyre, which exerts a major control on nutrient
availability in surface waters along the North American continental margin
(Pelegrí et al., 1996). Over the South Atlantic Bight region,
encompassing the Blake Plateau, the key mechanism fueling modern
phytoplankton production is sub-mesoscale frontal eddies arising from
meanders on the landward side of the Gulf Stream System (GSS; Richardson,
2001; Gula et al., 2015). Comparable cyclonic eddies of <100 km
diameter are also observed in other western boundary current (WBC) systems,
which are generally viewed as oligotrophic settings (Roughan et al., 2017).
These eddies are responsible for upward pumping of nutrients from
sub-thermocline, nitrate-rich waters (Lee et al., 1991). Upwelled waters
intrude onto the continental margin and sustain rich biological production
through the lifespan of an eddy (Roughan et al., 2017). Siliceous plankton
production and export in the GSS is influenced by a number of factors,
including Atlantic Meridional Overturning Circulation (AMOC) intensity and
the North Atlantic Oscillation, which together act to shift the GSS position
relative to the North American seaboard on a decadal timescale
(Sanchez-Franks and Zhang, 2015). Also, the topography of the North American
continental margin (Richardson, 2001) in conjunction with eustatic sea level
variations exert a strong influence on the GSS path on long timescales, with
features such as the Charleston Bump acting to deflect the jet trajectory
toward the open ocean (Pinet et al., 1981; Gula et al., 2015).
A northeastward-flowing, wind- and Coriolis-force-driven WBC likely operated
in the North Atlantic at least since the Cretaceous (Gradstein and Sheridan,
1983), albeit at reduced strength relative to the modern era before the final
closure of the Central American Seaway (Montes et al., 2012). Given the
overall stability of the western North Atlantic topography over the
Cenozoic, cyclonic frontal eddies were likely an inherent
feature of the South Atlantic Bight region throughout the Paleocene and
Eocene. The semi-periodic fluctuations in BN bioSiO2 flux through
time could therefore also be attributed to changes either in the mean GSS
path (e.g., Wade and Kroon, 2002) or variations in sub-thermocline nutrient
supply, which are largely dependent on vertical mixing of the ocean (Miskell
et al., 1985; Moore et al., 2008) – or a combination of both processes.
Reconstructing intermediate-water and deepwater circulation patterns in the
North Atlantic through the early Cenozoic is more complex than
reconstructing GSS history. Vahlenkamp et al. (2018) reviewed the existing
perspectives on the Atlantic Ocean circulation through the Paleogene.
εNd reconstructions generally indicate a southern
high-latitude source for the deep waters bathing the North American margin
throughout the early to middle Paleogene (Thomas et al., 2003; Batenburg et
al., 2018), although a Tethyan-sourced water mass is also hypothesized by
some researchers (Fontorbe et al., 2016; Vahlenkamp et al., 2018). At present,
it is not known how a southern-sourced, northward-flowing deepwater mass
may have affected nutrient availability and upwelling in the western North
Atlantic, especially along continental margins. The high diatom : radiolarian
(D:R) ratios (Witkowski et al., 2020b, for further discussion see below) and common occurrence of well-preserved epiphytic diatoms such as
Arachnoidiscus (see Witkowski et al., 2020a) suggest that much of the BN bioSiO2 flux through the Paleocene may be attributed to neritic production.
Varying proportions of continental-runoff-derived vs. upwelled nutrient
input could also be invoked to explain the disparity in bioSiO2
fluxes between the more proximal Site 1051 and the more distal Site 1050
through the Paleocene.
An intensely debated question in the early Paleogene deepwater circulation
reconstructions is the timing of the onset of NCW flow – a precursor to
quasi-modern deepwater circulation (Via and Thomas, 2006). North Atlantic
δ13C records do not indicate major paleocirculation changes
prior to the late Eocene (∼38 Ma; Katz et al., 2011; Borrelli
et al., 2014; Coxall et al., 2018), and numerous studies place the onset of
AMOC either shortly prior to or following the EOT (Via and Thomas, 2006;
Abelson and Erez, 2017; Coxall et al., 2018).
In contrast to the timing of NCW flow initiation indicated by isotopic proxy
records, the onset of widespread drift deposition in the North Atlantic is
documented considerably earlier, i.e., near the termination of the EECO
(∼49 Ma; Hohbein et al., 2012; Boyle et al., 2017). This is
also synchronous with ubiquitous deep-sea erosion coincident with the EMET
(Aubry, 1995; Witkowski et al., 2020b), strongly suggesting that the onset
of vigorous northern-sourced bottom current activity began at
∼49–47 Ma (Vahlenkamp et al., 2018; Witkowski et al., 2020a).
Following the EMET, the northward-flowing GSS and the invigorated deep WBC
facilitated diapycnal mixing, which likely enhanced biological pump
efficiency along continental margins of the western North Atlantic. This is
consistent with a range of geochemical proxies, including thallium isotope
(ε205Tl) evidence for increased Corg burial from
∼50 Ma (Nielsen et al., 2009) and with surface-to-deep
δ13C gradients (Hilting et al., 2008). BN diatom assemblage
data from Witkowski et al. (2020b) also support an oligotrophic regime over
BN for the time period prior to and including the EECO based on high
percentages of hemiauloids. Following the EECO (after ∼49 Ma), elevated percentages of diatom resting spores point to alternating,
perhaps seasonal, periods of nutrient enrichment and depletion, in line with
strong periodic upwelling of nutrients by means of Gulf Stream frontal
eddies (Lee et al., 1991). This interpreted invigoration in ocean mixing led
to a considerable increase in primary production during the early middle
Eocene, as evidenced by a rapid increase in both CaCO3 and
bioSiO2 fluxes at BN at ∼46 Ma (Fig. 4a, b).
Eocene biogenic silica (a) and calcium carbonate (b) fluxes at Blake Nose sites plotted against biogenic silica (c) and calcium carbonate (d) fluxes at eastern equatorial Pacific sites. Blake Nose carbonate data are from Shipboard Scientific Party (1998c, d, f). Eastern equatorial Pacific data are from Moore et al. (2008) and Lyle et al. (2005). Abbreviations: bioSiO2 – biogenic silica; GTS2012 – Geologic Time Scale 2012 (see Gradstein et al., 2012); MLET – Middle–Late Eocene Turnover; LLTM – Late Lutetian Thermal Maximum; ESAE – Eocene silica accumulation event; CAE – carbonate accumulation event.
Comparison to eastern equatorial Pacific bioSiO2 flux records
The only published early Paleogene bioSiO2 flux record from
another region of comparable duration to the BN composite record is that
derived from eastern equatorial Pacific (EEP) cores (Moore et al., 2008).
There are several important differences between the Atlantic and Pacific
records (Fig. 4), including (1) contrasting proportions of diatoms in the BN
vs. EEP sediments, (2) the presence or absence of exported neritic material,
and (3) a shorter time interval covered by the EEP record (Eocene only).
Where the records overlap in time, however, there is a variable degree of
coupling between bioSiO2 flux records from BN and EEP.
High diatom : radiolarian (D:R) ratios in early to middle Paleogene sediments at BN were interpreted by Witkowski et al. (2020b) to indicate that
preserved bioSiO2 was mostly of diatom origin. It is challenging,
however, to provide a quantitative estimate of the diatom vs. radiolarian
contribution to total bioSiO2 at BN, primarily because no
published diatom valve weight data are available. In the modern oceans,
radiolarian tests are on average an order of magnitude heavier than diatom
valves (with differences in fact ranging over several orders of magnitude;
Lisitzin, 1971). Assuming a 10:1 radiolarian-to-diatom skeleton weight ratio for the early Paleogene and using an average radiolarian test weight of 0.225 µg consistent with the range of values displayed by the oldest materials included in Moore (1969), we make a rough estimate of the diatom
vs. radiolarian contribution to total BN bioSiO2 based on quantitative siliceous microfossil counts of Witkowski et al. (2020b). As
other siliceous plankton groups are sparse in BN siliceous microfossil
assemblages (Witkowski et al., 2020b) and likely contribute little
bioSiO2 to sediments (Lisitzin, 1971), we exclude the relatively
minor contributions of silicoflagellates, siliceous dinoflagellates, and
chrysophycean cysts from the calculations. These rough approximations
indicate a mean diatom contribution of ∼35.7 % to the total
biogenic silica content at BN (Fig. S6), with the highest values observed for the early and middle Paleocene, consistent with the D:R ratios ranging as high as >200 reported by Witkowski et al. (2020b). Given that diatom valves are less resistant to dissolution than radiolarian tests, the contribution of diatoms to total bioSiO2 in early Paleogene sediments at BN is likely underestimated due to selective dissolution.
It is important to note that these considerations disregard the contribution
of siliceous sponge spicules to total bioSiO2 in the BN sediments. Counting and identifying sponge spicules were beyond the scope of the present study, and, to the best of our knowledge, no quantitative
studies on sponge spicules from the BN cores have been performed thus far.
Consequently, we were not able to use published data to estimate sponge
spicule contribution to total bioSiO2 in early Paleogene sediments
at BN. Several recent studies point to a declining contribution of sponges to
the total bioSiO2 flux that is probably linked to diatom expansion in the late Mesozoic (Maldonado et al., 1999; Conley et al., 2017). Modern bioSiO2 flux attributed to sponges ranges from 25 to
48 Tg Si yr-1 and is an order of magnitude lower than the total bioSiO2 flux estimate for continental margins (140–235 Tg Si yr-1) and for the
deep sea (153 Tg Si yr-1) (Hayes et al., 2020). Sponge spicules are therefore
unlikely to have made a significant contribution to total bioSiO2
flux at BN through the early to middle Paleogene. SiO2 preservation is
again a related issue, as sponge spicules undergo dissolution at slower
rates compared to siliceous plankton valves or tests (Bertolino et al., 2017). Although some attempts have been made (e.g., Warnock and Scherer,
2015), there is currently no standardized quantitative measure of diatom
preservation in sediments, and the basic indicators of silica dissolution
are chert–porcelanite and clinoptilolite occurrences. As indicated above,
both chert–porcelanite and clinoptilolite occur only at isolated, narrow
levels at the sites included in bioSiO2 flux estimates in the
present study, and in some intervals diatom preservation can be
considered pristine. Hence, the assumption is that no extensive diatom silica
dissolution has occurred in Holes 1050A and C and 1051A, which would lead to
preferential preservation of the more dissolution-resistant sponge spicule
silica over the more dissolution-prone diatom and/or radiolarian silica.
Thus, our conclusion is that siliceous sponge spicules do not contribute
significantly to total bioSiO2 flux in the BN cores.
In contrast to the abundant presence of diatoms at BN, Moore et al. (2008)
refer to the near absence of diatoms in the Eocene EEP cores as an
“enigma”. Although Moore et al. (2008) do not specify whether or not this
observation is based on sieved residues (thereby potentially missing diatoms
in the smaller sediment fractions), other Paleogene diatom reports from
pelagic low-latitude Pacific sites corroborate this view (e.g., Fenner,
1984). We propose that the reason for this difference in diatom abundance in
sediments between BN and EEP is twofold. Firstly, most early Paleogene
diatom occurrences in the Atlantic are in marginal settings (Witkowski et
al., 2020b), where at least part of the preserved diatom assemblage may
originate from offshore export of neritic plankton, and diatom preservation
may be fostered by higher concentrations of Al (DeMaster, 2014; Hayes et
al., 2020). Secondly, the radiolarian-rich Pacific sites mostly represent
pelagic deposition at water depths of ∼4–5 km. Diatom dissolution is facilitated by longer times in transit through the water column and longer times resting on the seafloor in slowly accumulating pelagic settings. The BN and EEP records also differ in the magnitude of
bioSiO2 fluxes: as discussed in Witkowski et al. (2014, 2020b),
the BN area received large volumes of neritic plankton through the early
Paleogene, which underwent offshore export likely by means of frontal
eddies, and the resultant bioSiO2 fluxes are high. The pelagic EEP
sites likely record only local pelagic production and deposition, with low
bioSiO2 fluxes relative to the BN.
Eocene EEP bioSiO2 accumulation rates are generally low,
punctuated by a series of elevated flux events termed ESAEs (Eocene silica
accumulation events – see Moore et al., 2008) and CAEs (carbonate
accumulation events – see Lyle et al., 2005) (Fig. 4c–d). Between ∼55 and 46 Ma, bioSiO2 fluxes in the EEP are low and appear to be decoupled from the BN records. ESAE 3 at ∼45.8 Ma
marks the onset of enhanced bioSiO2 flux in the EEP (Fig. 4c).
Notably, ESAE 3 appears to be age-equivalent to a prominent increase in BN
bioSiO2 and CaCO3 fluxes (Fig. 4a vs. c). ESAE 4 at
∼44.3 Ma is correlative with the peak in middle Eocene bioSiO2 flux at BN (Fig. 4a vs. c). Following ESAE 4, however,
trends in bioSiO2 flux again become decoupled between the two
regions. bioSiO2 fluxes diminish at BN between ∼42 and 38 Ma, and a concomitant reduction is observed in the geographic
distribution of siliceous microfossils in the Atlantic Ocean (Witkowski et
al., 2020b) (Fig. 4a vs. c). Thus, reduced bioSiO2 accumulation
between ∼42 and 38 Ma is not a local phenomenon restricted to the BN area, but instead it is likely indicative of a major change in nutrient supply or paleocirculation that affected the entire Atlantic basin.
In contrast, this late middle Eocene period of low bioSiO2 flux at
BN precisely corresponds to an interval of elevated bioSiO2 fluxes in the EEP (Fig. 4a vs. c), including the bimodal ESAE 5, which represents
the peak in the Eocene EEP bioSiO2 accumulation at ∼41 Ma. Thus, the decrease in nutrient levels in the Atlantic appears to have been associated with nutrient enrichment and elevated biosiliceous production in the EEP, representing an inter-basin shift in biosiliceous productivity and sedimentation.
Near the end of the middle Eocene, ESAE 6 is abruptly terminated at
∼38.5 Ma, concomitant with a major radiolarian turnover at EEP sites (Moore et al., 2008). A similar episode of accelerated turnover in radiolarians has been identified at ∼38.25 at BN (Kamikuri
and Wade, 2012; Newsam et al., 2017), but in conjunction with a rapid rise
in bioSiO2 flux levels. ESAEs 7 and 8 in the EEP are minor events
at ∼35 and ∼34 Ma, respectively, and no age-equivalent events are observed in the BN bioSiO2 flux record. We interpret the decoupling between BN and EEP bioSiO2 flux records after ∼42 Ma as a series of inter-basin bioSiO2 accumulation shifts (Fig. 4a vs. c), likely associated with deepwater circulation changes affecting nutrient availability in surface waters, but likely also impacting the seabed preservation of bioSiO2 (Berger, 1970).
Diminished BN bioSiO2 fluxes indicate lower nutrient supply from
∼42 to 38 Ma, i.e., through the interval spanning both the Late
Lutetian Thermal Maximum (LLTM, ∼41.5 Ma; Westerhold et al., 2018b) and Middle Eocene Climatic Optimum (MECO, ∼40 Ma; Bohaty et al., 2009; Henehan et al., 2020). Modern field observations
indicate that diminished supply of nutrients to the GSS may result from
weakened AMOC. Witkowski et al. (2020b) demonstrate a reduced geographic
range of biosiliceous accumulation in the Atlantic between ∼42 and 38 Ma, which is also consistent with a diminished nutrient supply.
Accordingly, we propose a period of potentially weakened AMOC spanning both
the LLTM and MECO events. In the EEP, the high rates of primary production
are sustained by advection of nutrient-rich sub-thermocline waters
associated with the equatorial divergence (Fiedler et al., 1991). We propose
that the paleocirculation changes that led to the interpreted disruption to
AMOC may have manifested themselves by nutrient enrichment in the EEP.
Thus, alternating loci of biosiliceous sedimentation between the Atlantic
and Pacific during the middle Eocene likely resulted from circulation shifts
that exerted control over bioSiO2 production and burial.
The Pacific-to-Atlantic bioSiO2 flux shift at ∼38 Ma coincides with increased rates of radiolarian turnover and planktonic foraminiferal extinction designated as MLET by Kamikuri and Wade (2012).
Notably, the late Eocene is also believed to have been a period of pelagic
diatom proliferation, probably due to the radiation of holoplanktonic taxa
(Sims et al., 2006; Egan et al., 2013). Thus, MLET may have made a
significant impact on siliceous microplankton evolution and production
globally. Most importantly, however, the abrupt increase in
bioSiO2 flux at Site 1053 shortly after MLET took place in
conjunction with shifts in benthic foraminiferal δ13C (Katz et
al., 2011; Borrelli et al., 2014; Coxall et al., 2018) interpreted to mark
the onset of NCW export.
Thus, the repeated shifts in bioSiO2 fluxes between NW Atlantic
and EEP through the late middle and late Eocene suggest that the “lagoonal”
Atlantic (carbonate-burial-favoring) vs. “estuarine” Pacific (SiO2-burial-favoring) circulation pattern proposed by Berger (1970) was established in the lead-up to LLTM and temporarily reversed at the MLET
before its final re-establishment shortly before the EOT. These shifts were
likely driven by changes in deep-sea circulation patterns arising from both
tectonic evolution in the Northern and Southern Hemisphere (e.g.,
Norwegian–Greenland Sea and Drake Passage region, respectively) and
long-term Eocene climate change. Importantly, however, this interpretation
implies that the two scenarios for NCW export inception (at EOT: Borrelli et
al., 2014; Coxall et al., 2018; vs. at the end of the EECO: Hohbein
et al., 2012; Boyle et al., 2017; Vahlenkamp et al., 2018) are not
mutually exclusive. The patterns in Atlantic-to-Pacific bioSiO2 flux fractionation outlined above suggest a ∼4 Myr period of
the early AMOC disruption spanning the last of the Eocene greenhouse warming
events, i.e., LLTM and MECO. Moreover, the evidence supporting NCW inception
in the late Eocene or early Oligocene may in fact point to a re-invigoration
of AMOC flow following a period of weakened overturning circulation between
42 and 38 Ma.
Implications for the silicate weathering feedback operation mode
The silicate weathering feedback has been proposed as the key mechanism for
keeping the Earth surface temperatures within a habitable range over
105–106-year timescales (Walker et al., 1981; Kasting, 2019). By
consuming atmospheric CO2 and releasing alkalinity and dissolved
silicon (Penman, 2016), this feedback mechanism also influences key
biogeochemical cycles within the ocean–atmosphere system, resulting in a
tight coupling between the marine carbon and silicon cycles (Tréguer and
De La Rocha, 2013). In recent years, however, the operation of the silicate
weathering feedback through the Cenozoic has been disputed, with a special
focus on whether the strength of the link between climate and continental
weathering varies through time (Caves et al., 2016; van der Ploeg et al., 2018). One point of disagreement concerns the early Paleogene. In the
traditional view (hereafter “constant feedback strength scenario”), which
assumes a linear relationship between global temperature change and
weathering, the early Paleogene greenhouse climates should facilitate
increased rates of chemical weathering on land directly proportional to the
magnitude of CO2-driven warming (e.g., Misra and Froelich, 2012; Sluijs et al., 2013; Penman, 2016). An emerging alternative view (hereafter
“variable feedback strength scenario”) is that during the Eocene the
feedback strength was at a minimum level (Caves et al., 2016; van der Ploeg
et al., 2018), with lowered silicate weathering intensity (and, hence,
reduced weathering feedback strength) promoting high pCO2 levels and a warm climate (Misra and Froelich, 2012). Assuming that the early to middle Paleogene silicon cycle already operated in its present-day form (Fontorbe et al., 2016; Conley et al., 2017), these scenarios should lead to different marine bioSiO2 flux responses. In the constant feedback strength scenario, bioSiO2 production and burial would be expected to peak during the EECO, i.e., the warmest period of the Cenozoic era (e.g.,
Kirtland-Turner et al., 2014; Cramwinckel et al., 2018; Westerhold et al., 2018a). In the variable feedback strength scenario, the silicate weathering flux should decrease through the Eocene (Caves et al., 2016), leading to a decrease in dissolved silicon supply and bioSiO2 burial.
In the BN composite, both %bioSiO2 and bioSiO2 flux values are high in the lead-up to the EECO (∼55 through 53.5 Ma) and considerably lower in the final phases of the EECO (∼51 through 49 Ma). The hiatus spanning ∼53.5 through 52 Ma, however, precludes any definitive conclusions on the behavior
of the silicate weathering feedback through the entire EECO period,
particularly with regard to the constant vs. variable strength of its
link to climate. However, bioSiO2 flux values through the middle
Eocene are similar to or consistently higher than background Paleocene–early
Eocene values in both the Atlantic and the Pacific (see also Moore et
al., 2008, and Sect. 4.2), which cannot be easily reconciled with the
linear feedback strength scenario. Thus, it appears that long-term trends in
bioSiO2 flux are more consistent with the variable feedback
strength scenario, suggesting that the strength of the link between climate
and terrestrial silicate weathering may indeed be variable through time.
Secondly, the high levels of bioSiO2 flux through the middle
Eocene cooling (Fig. 2b) point to enhanced nutrient supply from invigorated
ocean circulation as a major control on bioSiO2 flux in the
younger part of our study period.
Conclusions
Deepwater temperatures, atmospheric greenhouse gas levels, and continental weathering are identified as the main drivers of bioSiO2 flux through the Paleocene and Eocene at Blake Nose in the western North Atlantic Ocean. Variations in bioSiO2 fluxes support an early export of NCW, but also suggest a period of disruption due to diminished AMOC between ∼42 and 38 Ma, as suggested by the Atlantic-to-Pacific bioSiO2 flux fractionation. NCW export likely became
re-invigorated in the late Eocene, as indicated by a pulse of bioSiO2 flux and published paleocirculation proxy records.
Additionally, BN bioSiO2 fluxes indicate that the long-term
behavior of the silicate weathering thermostat conforms to the variable
rather than constant weathering–climate feedback strength scenario.
Overall, this study also demonstrates that disentangling silicate weathering, productivity, and paleocirculation controls on bioSiO2 flux records is challenging. While the globally integrated flux of
bioSiO2 to sediments must respond to global weathering rates,
bioSiO2 flux records at individual sites do not necessarily
reflect changes in the global flux because of site-specific or regional
effects like circulation change. Our hope is that by continuing to develop
bioSiO2 flux records from different parts of the oceans a
comprehensive picture may emerge from future studies. These records will
necessarily need to be constructed in conjunction with other lines of
evidence (i) to constrain the ancient silicon cycle and changes in dissolved
silicate concentration through application of silicon isotope (δ30Si) proxies and other techniques, (ii) to assess paleocirculation
changes through approaches such as fish-tooth neodymium isotopes, and (iii) to construct more sophisticated age models with refined sedimentation rate estimates, for example, through application of cyclostratigraphic approaches to achieve resolution on astronomical timescales.
Code and data availability
All data generated in this study are included in the
Supplement.
The supplement related to this article is available online at: https://doi.org/10.5194/cp-17-1937-2021-supplement.
Author contributions
JW, SMB, and DEP designed the study. JW, KB, BSW, and EM
performed sampling and analyses. EM performed the statistical analysis. All
authors participated in interpreting the data. JW prepared the paper
with input from all co-authors.
Competing interests
The contact author has declared that neither they nor their co-authors have any competing interests.
Disclaimer
Publisher's note: Copernicus Publications remains neutral with regard to jurisdictional claims in published maps and institutional affiliations.
Acknowledgements
Jakub Witkowski acknowledges the support from the National Science Center (Poland). Walter Hale, Holger Kuhlmann, and the IODP Bremen
Core Repository staff are thanked for efficient handling of multiple sample
requests. Annette Olivarez Lyle, Mitchell Lyle (who also provided feedback
on an earlier version of this paper), and Dorota Burska are thanked for
advice on using the alkaline leaching method. We are indebted to Julita
Tomkowiak, Agnieszka Ławecka, Adrianna Szaruga, Adrianna Januszkiewicz, and
Zofia Stachowska for assistance in sample treatment and spectrophotometric
analyses. John Barron and Louisa Bradtmiller are thanked for their
constructive reviews.
Financial support
This research has been supported by Narodowe Centrum Nauki (grant no. 2014/13/B/ST10/02988).
Review statement
This paper was edited by David Thornalley and reviewed by John Barron and Louisa Bradtmiller.
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