Oceanic organic carbon as a possible first-order control on the carbon cycle during the Bathonian–Callovian

Journal Pre-proof Oceanic organic carbon as a possible first-order control on the carbon cycle during the Bathonian–Callovian

Ricardo L. Silva, Luís V. Duarte, Grant D. Wach, Natasha Morrison, Taylor Campbell PII:

S0921-8181(19)30543-0

DOI:

https://doi.org/10.1016/j.gloplacha.2019.103058

Reference:

GLOBAL 103058

To appear in:

Global and Planetary Change

Received date:

26 March 2019

Revised date:

5 October 2019

Accepted date:

14 October 2019

Please cite this article as: R.L. Silva, L.V. Duarte, G.D. Wach, et al., Oceanic organic carbon as a possible first-order control on the carbon cycle during the Bathonian–Callovian, Global and Planetary Change(2019), https://doi.org/10.1016/ j.gloplacha.2019.103058

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© 2019 Published by Elsevier.

Journal Pre-proof Oceanic organic carbon as a possible first-order control on the carbon cycle during the Bathonian–Callovian

Ricardo L. Silvaa,b,* [email protected], Luís V. Duartec, Grant D. Wacha, Natasha Morrisona, Taylor Campbella

Basin and Reservoir Lab, Department of Earth Sciences, Dalhousie University, Life Sciences

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iCRAG, Department of Geology, School of Natural Sciences, Trinity College Dublin, The

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b

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Center, 1355 Oxford Street, Halifax, Nova Scotia B3H 4R2, Canada

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University of Dublin, College Green, Dublin 2, Ireland

MARE - University of Coimbra, Department of Earth Sciences, Rua Silvio Lima, Polo II, 3030-

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*

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290, Coimbra, Portugal

Corresponding author at: Department of Geology, School of Natural Sciences, Trinity College

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Dublin, The University of Dublin, College Green, Dublin 2, Ireland.

Journal Pre-proof Abstract Oceans are the largest, readily exchangeable, superficial carbon reservoir; a current challenge in investigating past and present environments and predict future evolution relates to the role of oceanic carbon in regulating Earths’ carbon cycle and climate. At least one paired δ13Ccarb-TOC decoupling event is noted in the Late Bathonian–Early Callovian. Provokingly, we suggest that this decoupling event and other carbon isotopic events

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in the Bathonian–Callovian resulted from the addition and removal of carbon from the oceanic

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organic carbon (OOC) reservoir, likely dominated by dissolved and fine particulate oceanic

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organic carbon. The decoupling event is characterised by a mainly invariant δ13CTOC record and

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increasingly more positive δ13Ccarb values. The δ13Ccarb-TOC decoupling is tentatively explained by the expansion of the OOC reservoir, which increased the residence time of carbon in the oceans

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and buffered (slow equilibrium) δ13COOC (approximated by δ13CTOC) to changes in δ13C of

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oceanic inorganic carbon, comprised mainly of dissolved inorganic carbon (approximated by δ13Ccarb). Reconversion of OOC to CO2 may have resulted in negative δ13C excursions and

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increased atmospheric pCO2, whereas the change from accumulation of OOC to export of

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organic carbon to sediments may have resulted in positive δ13C excursions and widespread accumulation of organic matter. It is speculated that other small-scale (~ 1‰) δ13Ccarb-TOC excursions in the Mesozoic record may have resulted from similar episodes of OOC accumulation and oxidation. Our study calls attention to the potential of the OOC reservoir to act as a critical driver for planetary-scale changes over short geological time intervals. Keywords: δ13Ccarb-TOC; carbon cycle; palaeoceanography; Middle Jurassic; Lusitanian Basin; Portugal

Journal Pre-proof 1. Introduction The carbon cycle is usually divided into two cycles, the short- and long-term carbon cycles (Berner, 1999). Carbon in the short-term carbon cycle is rapidly exchanged within the surficial reservoirs, consisting of the oceans, atmosphere, biosphere, and soil, whereas carbon in the long-term cycle is slowly exchanged between the geosphere and the surficial systems (Berner, 1999; Ciais et al., 2013). The short-term carbon cycle is the dominant control on

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atmospheric CO2 over millennia, whereas the long-term carbon cycle controls atmospheric CO2

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and O2 over millions of years (Berner, 1999, 2006).

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Oceans are, by far, the largest readily exchangeable superficial carbon reservoir; a current challenge in investigating past and present environments and predict future climatic evolution

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concerns the role of the oceanic organic and inorganic carbon reservoirs in regulating Earths’

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carbon cycle and climate (Ciais et al., 2013; Hansell, 2013; Legendre et al., 2015). For example,

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recent models suggest that ocean warming and freshening of surface waters is causing a slowdown of overturning circulation, which is projected to continue for the next millennia

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(Schmittner et al., 2008; Rahmstorf et al., 2015). It is postulated that accumulation of organic

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carbon in the oceans might increase in these conditions, thus providing negative feedback to rising atmospheric CO2 (Jiao et al., 2010, 2014; Cavan et al., 2017; Shen and Benner, 2018).

1.1. Oceanic carbon in modern oceans Oceanic carbon is here divided into two large pools, oceanic inorganic carbon (OIC) and oceanic organic carbon (OOC), usually characterised as inorganic or organic or and then in dissolved or particulate (Fig. 1a). Dissolved inorganic carbon (DIC) consists of CO2* (=CO2 + H2CO3), HCO3−, CO32− in solution (Freeman and Hayes, 1992). Particulate inorganic carbon

Journal Pre-proof (PIC) results mainly from the precipitation of CaCO3 shells by coccolithophores, foraminifers, and pteropods, for example. On the other hand, OOC is operationally divided into dissolved organic carbon (DOC) and particulate organic carbon (POC), determined by the pore size of the filters (usually 0.7 μm or, less frequently 0.45 or 0.2 μm) used to separate “dissolved” from “particulate” material, respectively (Emerson and Hedges, 2008; Ciais et al., 2013; Carlson and Hansell, 2015) (Fig. 1b).

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With a residence time of about 100 kyr (Emerson and Hedges, 2008), DIC is the largest

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carbon reservoir in today’s oceans (~38,000 PgC) (Ciais et al., 2013). The marine biota (~3 PgC)

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represents only a very small fraction of the oceanic carbon inventory and have a short residence

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time (Ciais et al., 2013). In today’s oceans, recalcitrant DOC is the largest exchangeable carbon component of the OOC reservoir (Fig 1). The refractory DOC subclass holds about 630±32 PgC,

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similar to the size of the atmospheric carbon pool, consisting of ~832 PgC of which CO2 is the

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dominant carbon-bearing gas and includes CH4 (~3.7 PgC) and CO (~0.2 PgC) (Ciais et al., 2013; Hansell, 2013). Refractory DOC comprises about 95% of the recalcitrant DOC reservoir

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and has an estimated residence time of about 16 kyr, indicating that oceans can store organic

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carbon for extended periods of time and that geological variations in the size of the OOC reservoir, and particularly recalcitrant DOC, could impact atmospheric CO2 concentrations (Hansell, 2013) and, therefore, climate evolution and biological and geological processes. In this geological study, dissolved and particulate oceanic organic carbon (DPOOC) refers to organic carbon that is truly dissolved as well as suspended colloidal organic carbon (i.e. labile and recalcitrant DOC) and fine-grained POC (see discussions in (Ogawa and Tanoue, 2003; Verdugo et al., 2004; Fike et al., 2006; Emerson and Hedges, 2008; Swanson-Hysell et al.,

Journal Pre-proof 2010; Nebbioso and Piccolo, 2013) (Fig. 1b). The DPOOC component of the OOC reservoir is most likely dominated by recalcitrant DOC. Equilibrium and kinetic fractionation among organic matter (OM), CO2, and HCO3control the carbon isotope ratios of DIC, atmospheric CO2 and OOC (Emerson and Hedges, 2008). The main processes that affect the 13C/12C ratio of the “superficial” carbon reservoirs at geological time intervals are the equilibrium between DIC and atmospheric CO2, fractionation

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between DIC and carbonate minerals; fixation of dissolved CO2 and production of biomass via

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photosynthesis, respiration, remineralisation, preservation, and sedimentary reworking and

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resuspension of OM in the oceans and sediments (Hayes et al., 1999). There is little isotopic discrimination effect during respiration but, during photosynthesis 12C is taken up preferentially

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instead of 13C, leaving the superficial DIC pool enriched in 13C. As OM sinks into intermediate

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and deep waters, “sinking” organic carbon is remineralised and the deeper ocean DIC pool

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becomes enriched in 12C. At geological time scales, removal of isotopically “light” OM into the sedimentary organic matter (SOM) reservoir generates a relative enrichment in 13C of dissolved

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CO2 in seawater and continuation of this process leads to an overall increase in 13C of marine

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carbonates and OM (Jenkyns and Clayton, 1986; Wefer and Berger, 1991; Marshall, 1992; Patterson and Walter, 1994; Mitchell et al., 1996; Weissert et al., 1998; Holmden et al., 1998; Kump and Arthur, 1999; Hesselbo et al., 2000; Berner, 2003; Immenhauser et al., 2003; Swart and Eberli, 2005; Panchuk et al., 2006; Newton and Bottrell, 2007; Emerson and Hedges, 2008; Silva et al., 2011; Silva and Duarte, 2015).

1.2. Investigating oceanic carbon in the geological record

Journal Pre-proof Paired δ13Ccarb-TOC (δ13C (‰) = [(13C/12Csample)/(13C/12Cstandard)–1]*1000) analysis is one of the main tools to investigate perturbations of the carbon cycle at geological scales, where changes in δ13Ccarb and δ13CTOC are taken as proxies to changes in the 13C/12C ratio of OIC (dominated by DIC) and OOC (dominated by DPDOC, see above) respectively (Freeman and Hayes, 1992; Hayes et al., 1999; Berner, 2003) (Fig 1a). A conventional way to investigate the carbon cycle at geological time scales assumes that the carbon cycle is in (or near) steady state.

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According to Krissansen-Totton et al. (2015), this assumption specifically means that i) the

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carbon input into surficial reservoirs (atmosphere, oceans, and biosphere) is balanced by the

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burial of carbonates plus organic carbon, and that ii) 13C/12C of inputs equal the average 13C/12C

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of outputs (Schidlowski et al., 1979; Berner et al., 1983; Garrels and Lerman, 1984; Schidlowski, 1988; Beerling and Berner, 2002; Wallmann and Aloisi, 2012; Krissansen-Totton et al., 2015).

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This approach usually groups the atmosphere, oceans, life and soils, thus avoiding having to

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explicitly include the short-term carbon cycle (Berner, 2004). This is a reasonable approximation for time intervals in the order of millions of years, but it may not hold when analysing high-

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resolution (hundreds to hundreds of thousands of years) sedimentary records (Sundquist, 1991;

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Berner, 2004). At short-time scales, redistribution of carbon among the superficial reservoirs, likely driven by earth-orbital variations and internal ocean-atmosphere-biosphere-cryosphere feedbacks, may lead to “non-steady state” effects in the carbon cycle (Sundquist, 1985, 1991). For example, the observation of a temporal trend in the last 50 years in the Redfield ratio (carbon:nitrogen:phosphorus) of particle flux to the deep ocean suggests that the biological portion of the marine carbon cycle is not in steady state (Pahlow and Riebesell, 2000). In the geological record, Rothman et al. (2003) observed decoupling of the δ13Ccarb-TOC record during the Neoproterozoic. These authors suggested that in discrete periods of the

Journal Pre-proof Neoproterozoic, the carbon cycle was in non-steady state due to the accumulation of a large OOC pool (assumed to be at least 10x the size of the inorganic reservoir) and that the longer residence time of carbon in this reservoir buffered its isotopic response (i.e. slow equilibrium) to changes in δ13C of DIC. Rothman et al. (2003) also discussed that DOC may have dominated the OOC pool and that the large OOC reservoir lead to fluctuations in the isotopic composition of the smaller DIC reservoir either passively, by changing the isotopic composition of the fluxes

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between DIC and OOC (ε0 not constant) or actively, due to enhanced organic remineralisation.

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More recently, it was proposed that the accumulation and remineralisation of oceanic

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DOC may have modulated the carbon cycle during the Paleocene and Eocene (Sexton et al.,

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2011). Several intervals of rapid and pronounced global warming are recognised in this interval, including the Paleocene-Eocene Thermal Maximum (PETM). It was suggested that the less

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severe hyperthermals after the PETM resulted from astronomically paced large-scale releases of

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DOC previously accumulated in anoxic areas due to ocean ventilation (Sexton et al., 2011). It was posited that the small isotopic fluctuations of about 1 ‰ associated with possible

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hyperthermals after the PETM resulted from the periodic accumulation and subsequent oxidation

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of about 1600 Pg C from the DOC reservoir, approximately 2x the size of the DOC reservoir in today’s oceans (Hansell, 2013). Despite the uncertainty regarding, for example, the unknown sensitivity of refractory DOC degradation to ocean oxygenation, mechanisms for DOC accumulation, and oxidant availability (Bristow and Kennedy, 2008; Johnston et al., 2012; Ridgwell and Arndt, 2015), the studies of Rothman et al. (2003) and Sexton et al. (2011) call attention to the possibility of a nonsteady state carbon cycle and the possibility of OOC and DPOOC to act as a driver for planetaryscale changes at geological time scales.

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1.3. Oceanic carbon as a driver for planetary-scale changes in the Mesozoic? The Bathonian–Callovian (168–163 Ma; Ogg et al., 2016), archives major and recurring sedimentary, biological, and climatic planetary-scale changes, including the hypothesised formation of polar ice, demise of carbonate platforms, eutrophication, high environmental stress, and several perturbations of the carbon cycle (Kenig et al., 1994; Hubbard and Boulter, 1997;

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Price, 1999; Jenkyns et al., 2002; Dromart et al., 2003b, 2003a; Guex, 2006; O’Dogherty et al.,

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2006; Nunn et al., 2009; Andrieu et al., 2016; Koevoets et al., 2016). However, the drivers

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behind these planetary-scale changes and its consequences are still poorly understood (Dromart

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et al., 2003b; Jenkyns et al., 2012; Arabas, 2016; Ikeda et al., 2016; Alberti et al., 2017; Song et al., 2017).

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We report here a paired δ13Ccarb-TOC record from the Upper Bathonian–Lower Callovian

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hemipelagic succession cropping out at Praia da Pedra do Costado, Lusitanian Basin (Portugal). The δ13Ccarb-TOC record for this section is characterised by a short-lived decoupling event in the

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uppermost Bathonian–lowermost Callovian. The decoupling between δ13Ccarb-TOC suggests that

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the short-term carbon cycle may have played a significant role in this event. In this paper, we examine potential causes for the δ13Ccarb-TOC decoupling event and their significance to the understanding of the oceanographic and environmental drivers that modulated the carbon cycle and marine dynamics in the Bathonian–Callovian.

2. Geological Background The Lusitanian Basin is a narrow and small North–South elongated basin, located in the western Iberian Margin (Wilson et al., 1989; Pinheiro et al., 1996; Rasmussen et al., 1998; Alves

Journal Pre-proof et al., 2002) (Fig. 2). The Lower to Middle Jurassic interval belongs to the first 1st-order sedimentary cycle of the Lusitanian Basin (Wilson et al., 1989; Alves et al., 2002), registered mostly by carbonate and shale deposits (Soares et al., 1993; Azerêdo et al., 2003). The temporal and lateral facies variation of theses successions highlight the basin evolution from a homoclinal ramp setting during the Lower Jurassic (e.g. Duarte et al., 2007; Silva et al., 2015) to a distally steepened ramp during the Middle Jurassic (Azerêdo, 1998; Azerêdo et al., 2002, 2014, 2020)

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(Fig. 2). As the result of the regressive phase (started in the Lower Toarcian) of the 1st-order

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sedimentary cycle, two main sedimentary domains are differentiated in the Lusitanian Basin

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during the Middle Jurassic: to the east, the sedimentary series are typified by inner carbonate

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ramp facies; to the western sectors, carbonate hemipelagic deposits crop out at Baleal and Cabo Mondego (Mangold, 1990; Azerêdo, 1998; Azerêdo et al., 2002, 2003, 2014; Canales and

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Henriques, 2008; Fernández-López et al., 2009) (Figs. 2 and 3).

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The Cabo Mondego section, corresponding to the coastal areas of Cabo Mondego (Figueira da Foz, Portugal) contains the informal Cabo Mondego formation. This section is a

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worldwide stratigraphic reference that includes the Bajocian Global Boundary Stratotype Section

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and Point (Pavia and Enay, 1997) and the Bathonian Auxiliary Section and Point (FernándezLópez et al., 2009). In the Cabo Mondego region, the Cabo Mondego formation encompasses the entire Middle Jurassic (e.g., Cariou et al., 1988; Mangold, 1990; Azerêdo et al., 2003; Fernández-López et al., 2009; Henriques et al., 2016) (Fig. 3). The Bathonian–Callovian boundary interval is particularly expressive at this location and is the only point in the basin where this boundary has some biostratigraphic control based on ammonites (Cariou et al., 1988; Mangold, 1990). Another point of interest is the occurrence of centimetric organic-rich shales dated from the Upper Bathonian, interpreted to represent the onset of a deepening phase and

Journal Pre-proof sedimentary starvation (Soares et al., 1993; Azerêdo et al., 2003) (Fig. 3). The Bathonian– Callovian deposits are interpreted as a 2nd-order transgressive-regressive facies cycle (Azerêdo et al., 2014). The studied portion of the Cabo Mondego section is located at Praia (beach) da Lage do Costado (29T; 507777.89 m E; 4448435.05 m N), Cabo Mondego, Figueira da Foz, Portugal (Fig. 2–4). The sampled and studied interval is dated from the Upper Bathonian–Lower

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Callovian (Ruget-Perrot, 1961; Cariou et al., 1988; Mangold, 1990; Soares et al., 1993) and

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consist of about 40 metres of centimetric–decimetric shale(marl)–limestone alternations with a

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few centimetric apparently organic-rich beds exposed in a restricted wave-cut platform.

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Ammonite macrofauna allows for a good biostratigraphic control of the Bathonian series. It was suggested (Cariou et al., 1988; Mangold, 1990) that the Bathonian–Callovian limit is located 8–

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13 metres above beds 152/153 (bed FN9 in Mangold, 1990). In our measured section, this

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interval crudely corresponds to beds 180–210 (Fig. 4). The limit between the Middle–Upper Bathonian is poorly defined. Previous work

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(Mangold, 1990) have defined the Julii, Histricoides, and Angulicostatum subzones of the

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Retrocostatum Zone in the Cabo Mondego section. According to (Mangold, 1990), the Retrocostatum Zone corresponds to the entire Upper Bathonian. The Julii Subzone corresponds to the lower and middle portion of the Orbis Zone, the Histricoides subzone to the upper Orbis Zone, and the Angulicostatum subzone to the uppermost Orbis Zone and Discus zone as defined by Torrens in (Cope et al., 1980) for the NW Europe. The remaining of the Lower Callovian section was inaccessible during the sampling campaign (see also Mangold, 1990), covered by sand, boulders, and cliff debris. The cliff section presented obvious signs of weathering (e.g. discolouration, karstification) and was avoided from sampling.

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3. Materials and methods Seventy-four (74) samples were collected from the Upper Bathonian–Lower Callovian interval of the Cabo Mondego Formation in the Praia da Lage do Costado (Figs 2–4 and Table 1). From each sample, about 1-2 g of material was inspected for final cleaning, crushed, and then sent to Iso-Analytical Laboratories (United Kingdom) for determination of total organic carbon

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(TOC), δ13C and δ18O in bulk carbonates (carbonates), and carbonate-free (13CTOC) isolates

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(Table 1).

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TOC was calculated first by obtaining % Cwhole-rock from the weight loss of the

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decarbonated sample. The carbon content (total ion beam data) of the decarbonated sample (% Cdecarbonated) was measured on an elemental analyser (Europa Scientific Ltd) using IA-R001 wheat

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flour (40.2 %C w/w, Iso-Analytical Ltd, Crewe, UK) as the working reference material. The %

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Cdecarbonated data were then corrected for the weight that was lost during the decarbonation procedure to calculate TOC (%). This method was validated against commercial TOC analysers

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(Iso-Analytical Ltd un-published data).

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Stable carbon isotopes were determined in carbonate (13Ccarb) and carbonate-free (13CTOC) fractions. Stable oxygen isotopes were determined in carbonate (18Ocarb). All isotopic ratios are expressed relatively to VPDB (Coplen, 1994). The carbonate fractions were analysed by Continuous Flow–Isotope Ratio Mass Spectrometry (CF-IRMS). The samples were weighed into clean ExetainerTM tubes and then flushed with 99.995 % helium. After flushing, phosphoric acid was added to the samples, they were heated at 90ºC for 3 hours and then left to react in the acid overnight to allow complete conversion of carbonate to CO2. Reference and control materials were prepared the same way. CO2 was sampled into a continuously flowing He stream

Journal Pre-proof using a double holed needle and resolved on a packed column gas chromatograph and the resultant chromatographic peak carried forward into the ion source of a Europa Scientific 20-20 IRMS where it is ionized and accelerated. Gas species of different mass are separated in a magnetic field then simultaneously measured using a Faraday cup collector array (CO2 at m/z 44, 45, and 46). The reference materials used during analysis were IA-R022 (Iso-Analytical working standard calcium carbonate, 13C = -28.63 ‰ and 18O = -22.69 ‰), NBS-18 (carbonatite, 13C

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= -5.01 ‰ and 18O = -23.2 ‰) and IA-R066 (chalk, 13C = +2.33 ‰ and 18O = -1.52 ‰). IA-

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R022 was calibrated against and is traceable to, NBS-18 and NBS-19 (limestone, δ13C = +1.95

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‰). IA-R066 was calibrated against and is traceable to NBS-18 and IAEA-CO-1 (Carrara

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marble, δ13C = +2.5 ‰). NBS-18, NBS-19 and IAEA-CO-1 are inter-laboratory comparison standard materials distributed by the International Atomic Energy Agency (IAEA).

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For carbonate-free (13CTOC) isolates and Elemental Analysis- Isotope-Ratio Mass

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Spectrometry (EA-IRMS) analysis, weighed sub-samples were taken from the sample vials, placed in universal tubes, acidified with 2M hydrochloric acid, mixed, oven heated at 60 °C for 2

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hours and left for 24 hours to allow all carbonate to be liberated as CO2. The sample fractions

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were then isolated by centrifugation, and the acid was then decanted. The samples were then washed twice using distilled water and underwent centrifugation. After acid washing, the fractions were oven dried at 60 °C. Tin capsules containing a sample or reference material are loaded into an auto-sampler on a Europa Scientific elemental analyser, furnace held at 1000 °C and combusted in an oxygen-rich environment. The exhaust gas is reacted in a series of catalysts (e.g. chromate, silver wool) to oxidise hydrocarbons and remove sulphur and halides, purifying the CO2 stream. After purification, the remaining gas stream enters a Europa Scientific 20-20 IRMS ion source where it is ionised and accelerated. Gas species of different mass are separated

Journal Pre-proof in a magnetic field then simultaneously measured using a Faraday cup collector array (CO2 at m/z 44, 45, and 46). The reference material used during δ13CTOC analysis was IA-R001 (wheat flour, 13CV-PDB = of -26.43 ‰). For quality control purposes check samples of IA-R001, IAR005 (beet sugar, 13CV-PDB = -26.03 ‰) and IA-R006 (cane sugar, 13CV-PDB = -11.64 ‰) were analysed during batch analysis of the samples. IA-R001, IA-R005 and IA-R006 are calibrated against and traceable to IAEA-CH-6 (sucrose, 13CV-PDB = -10.43 ‰). IAEA-CH-6 is an inter-

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laboratory comparison standard distributed by the International Atomic Energy Agency (IAEA).

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To ensure the precision of laboratory analyses, duplicated samples were introduced every

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five samples. Based on unpublished analyses and the data presented here, analytical precision is

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better than ±0.1‰.

4. Results

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TOC in the studied section varies between 0.2–4.5 % (Fig 4 and Table 1). Most of the determined values are below 1 %. The highest TOC values are observed in black shales dated

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from the Histricoides Subzone, Retrocostatum Zone (Mangold, 1990).

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δ13Ccarb varies from -0.45 to 3.16 ‰ (Fig 4 and Table 1). Overall, δ13Ccarb presents a slight negative trend for the ~10 m of the studied section dated from the Middle?-Upper Bathonian (Julii Subzone; Retrocostatum Zone (Mangold, 1990), followed by a positive trend, starting at the Julii–Histricoides subzone boundary through to the Lower Callovian (Mangold, 1990). Values of δ18Ocarb vary between -3.24 and -1.95 ‰ (Fig 4 and Table 1). Two δ18Ocarb maxima are observed; the first at the transition between the Julii and Histricoides subzones (Retrocostatum Zone) and the second at the Bathonian (Angulicostatum Subzone?)–Callovian

Journal Pre-proof boundary (Mangold, 1990). No significant correlation is observed between δ13Ccarb and δ18Ocarb (r2=0.16) (Fig 5). δ13CTOC varies from -27.39 to -23.18 ‰ (Fig 4 and Table 1). For the Bathonian, most of the δ13CTOC analyses range between around -27.3 to -26.5 ‰, decoupled from δ13Ccarb in the interval between the upper Bathonian (Retrocostatum Zone, Histricoides Subzone)–Lower Callovian (Mangold, 1990). The Lower Callovian succession records a sharp positive carbon

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isotopic excursion (CIE) of about 1.3 ‰. Positive δ13CTOC values are also observed around the

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black shale intervals (Retrocostatum Zone, Histricoides Subzone (Mangold, 1990)), around 30 m

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from the base of the studied section. These are the most positive values of the studied section, ranging between -24 and -23 ‰. No significant correlation is observed between TOC and

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δ13CTOC (r2=0.4252 considering the entire dataset, r2=0.172 without the high TOC (> 4 %)

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sample.

5. Diagenesis and preservation of a primary signal

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5.1. Are the δ13C and δ18O records of bulk carbonate a primary or diagenetic signal?

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The effects of early marine diagenesis on carbonate δ13C and δ18O are not yet fully understood (Swart, 2015). However, and as a rule of thumb, inorganic aragonite and high-Mg calcite cements tend to have more positive δ13C than biologically derived skeletal aragonite (Braithwaite and Camoin, 2011). Early cementation influenced by-products linked with microbial OM decay, resulting in the production of CO2 depleted in 13C, may result in lower δ13C values (Marshall, 1992; Swart and Melim, 2000; Dickson et al., 2008; Coimbra et al., 2009; Silva et al., 2011; Duarte et al., 2014).

Journal Pre-proof The supply of oxidants controls the extent of the degradation of OM in sediments and it is apparent that under normal marine conditions oxidation of OM does not impact δ13C significantly in carbonate-rich rocks (Swart, 2015). On the other hand, the occurrence of diagenetic components influenced by OM oxidation might be significant in low or nondeposition intervals and where carbonates are a minor component (Swart, 2015). δ18O values of this kind of cements are also likely to vary depending on the nature of the original grains and the

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temperature of precipitation. Precipitation of cement in cooler environments would result in

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more positive δ18O (Swart and Melim, 2000). Upon deeper burial, the temperature effect is

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further recorded on the carbonate δ18O, leading to lighter values due to the higher temperatures.

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δ13C is more resistant to change during late-stage diagenesis as it has a very low thermal dependency and requires a high degree of water-rock interaction (Marshall, 1992; Jacobsen and

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Kaufman, 1999). However, the presence of high pCO2 and low δ13C fluids, a situation not

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uncommon during late-stage diagenesis, can alter carbonate δ13C towards lighter values (Derry, 2010).

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The relatively carbonate-rich section and the low TOC contents of the studied samples

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(generally <0.5 %) suggests that the δ13Ccarb trend reflects the primary oceanographic signal, even if the absolute values are slightly modified by diagenesis (Figs 1, 4, and 5). The lack of correlation between δ13Ccarb and δ18Ocarb supports the interpretation of the absence of a strong diagenetic imprint in the observed δ13Ccarb trend (Fig. 5). The correlation of our δ13Ccarb record with several contemporaneous datasets indicates that the observed temporal trend is, at least, a regional feature (Fig. 6).

5.2. Organic matter source and significance of the δ13CTOC profile from the Lusitanian Basin

Journal Pre-proof The obtained δ13CTOC record represents a mixed δ13C signal derived from primary production by aquatic organisms, microbial biomass (including dark carbon), and terrestrial material, all superimposed by diagenesis, thus resulting in potentially highly variable SOC δ13C signatures (Tyson, 1995; Emerson and Hedges, 2008; Herndl and Reinthaler, 2013; Hoefs, 2015; Suan et al., 2015) (Figs 1 and 4). Carbon isotopic composition of OM mainly depends on i) 13C content of the carbon source, including mixing; ii) isotope effects associated with the

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assimilation of carbon; iii) isotope effects associated with metabolism and biosynthesis; and iv)

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cellular carbon budgets (Hayes et al., 1999; Hayes, 2001; Hoefs, 2015). Additional factors such

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as temperature, availability of CO2(aq), and nutrient availability were shown to exert some control on the δ13C of phytoplankton (Emerson and Hedges, 2008). Transformation of OM during its

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transit in the water column seems not to impact significantly δ13C of organic carbon in

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sediments. However, chemotrophic addition to OM deposited in modern oxygen minimum zones

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is shown to result in variably 13C depleted SOM (Rush et al., 2019) and the contribution of

(Galimov, 2006).

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reworked material makes SOM isotopically lighter (within 1–2‰) than its biological source

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After deposition and during early diagenesis, two main processes change the isotopic composition of SOC: i) selected preservation of lipids (13C depleted) in relation to the more easily degradable carbohydrates and proteins, and ii) isotope fractionation due to the metabolism of microorganisms (Tyson, 1995; Hayes, 2001; Galimov, 2006; Hoefs, 2015). Preservation of lipid-rich 13C depleted marine SOM under oxygen-depleted conditions may produce negative δ13CTOC anomalies that do not necessarily reflect changes in the δ13C of the DIC pool (Suan et al., 2015). On the other hand, high-temperature thermal maturation processes and isotope exchange reactions between kerogen and carbonates may lead to a 13C enrichment of kerogen

Journal Pre-proof (Chung and Sackett, 1979; Lewan, 1983; Peters, 1986; Clayton, 1991; Hoefs, 2015), resulting in a shift of δ13CTOC to heavier values. From the above, it is irrefutable that a complex array of biological and diagenetic processes control δ13CTOC from any study. Despite its limitations, and keeping in mind that virtually all δ13CTOC records likely deviate from original δ13C, the overall shape and rates of change of δ13CTOC profiles are still of great utility to trace changes in paleoenvironmental and

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paleoceanographic conditions contemporaneous with deposition (Oehlert and Swart, 2014; Jarvis

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et al., 2015; Suan et al., 2015).

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In well-preserved samples, the variation of OM composition is one of the main causes of δ13CTOC variation. We address this issue mainly by comparing the obtained δ13CTOC from

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Portugal with published δ13C data from organic marine and terrestrial material. Upper Bajocian–

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Bathonian fossil wood from Yorkshire (UK) presents an average δ13C around -23 to -22 ‰ and

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ranges between -24.2 to -21.6 ‰ (Hesselbo et al., 2003). In the mid-Callovian Oxford Clay Formation (UK), eleven (11) fossil wood samples have an average δ13C of -23.8 ‰ (Kenig et al.,

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1994). Kenig et al. (1994) estimated an average δ13C of -23.5 ‰ for terrestrial OM (δ13Cterrestrial)

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and between -29.1 to -26.6 ‰ for primary marine OM in the Oxford Clay Formation. Based on the studies of Kenig et al. (1994) and Hesselbo et al. (2003), we assume a δ13Cterrestrial between 23.5 and -22.5 ‰ for the Bathonian–Callovian (nonetheless, further studies are needed to evaluate short-term δ13Cterrestrial variation for this time interval). The obtained δ13CTOC from the Upper Bathonian–Lower Callovian in the Lusitanian Basin is invariant for the most part, ranging between about -27.0 to -26.5 ‰ (Figs 4 and 6). The difference of 3–4 ‰ between Bathonian–Callovian δ13Cterrestrial and δ13CTOC from Portugal suggests a limited terrestrial contribution in the δ13CTOC record from the Lusitanian Basin in the

Journal Pre-proof decoupling interval. This observation is supported by another Mesozoic example suggesting that typical marine SOM usually yields δ13CTOC values that are 3–4‰ lighter when compared to coeval fossil wood or SOM of terrestrial origin (Tyson, 1995; Suan et al., 2015). δ13Cterrestrial is commonly interpreted to reflect the carbon isotopic composition of atmospheric CO2 and, consequently, δ13CDIC via air-water gas exchange (Gröcke et al., 1999; Gröcke, 2002; Hesselbo et al., 2007). Therefore, if terrestrial carbon was a major contributor to the studied δ13CTOC record

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of the Lusitanian Basin, especially during the decoupling interval, the paired δ13Ccarb-TOC record

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should present some degree of co-variation. Finally, the invariant portion of the δ13CTOC curve in

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the decoupling interval (Figs 4 and 6) suggests a consistent and invariable isotopic composition

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of SOC and likely similar SOM assemblages (Tyson, 1995). Summarizing, the δ13CTOC record from the Lusitanian Basin in the decoupling interval is interpreted to dominantly reflect a marine

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signal.

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The interval between beds 140–166 (reaching up to about -24 ‰ in bed 164) and the last 5 m of the studied section (ranging between -26 to -25.5 ‰) present more positive δ13CTOC

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values. The first interval is dated from the Histricoides Subzone, Retrocostatum Zone, whereas

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the second is dated from the Lower Callovian (Mangold, 1990) (Fig. 4). As discussed above, and assuming a terrestrial OM enrichment in 13C of about 3–4‰ in relation to marine OM, we postulate that the high-TOC laminated shales may have a significant terrestrial contribution, thus explaining the distinction between high-TOC laminated shales with heavier δ13CTOC vs low-TOC marls with lighter, and largely, invariant δ13CTOC (Fig 4). The δ13CTOC positive excursion of about 1‰ at the base of the Callovian is interpreted to reflect a change in the 13C/12C ratio or composition of SOM. The broad significance of the Late Bathonian–Early Callovian decoupling event and the Early Callovian δ13CTOC positive CIE is discussed in the following sections.

Journal Pre-proof

5.3. The Bathonian–Callovian regional δ13C record The regional δ13Ccarb profile for the Bathonian–Callovian includes the Early–Late Bathonian δ13Ccarb negative CIE, Late Bathonian–Early Callovian δ13Ccarb positive CIE, and the Early Callovian δ13Ccarb negative CIE (Kenig et al., 1994; Bartolini et al., 1995, 1996; Jenkyns, 1996; Gruszczyński, 1998; Bartolini et al., 1999; Jenkyns et al., 2002; O’Dogherty et al., 2006;

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Nunn et al., 2009; Jach et al., 2014; Arabas, 2016; Koevoets et al., 2016; Alberti et al., 2017)

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(Fig. 6). In the well-dated La Cornicabra section (Spain), the Early–Late Bathonian δ13Ccarb

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negative CIE has an amplitude of about 0.5 ‰, the Late Bathonian–Early Callovian δ13Ccarb positive CIE has an amplitude of ~ 0.8 ‰, and the Early Callovian δ13Ccarb negative CIE has an

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amplitude of about 0.9 ‰ (O’Dogherty et al., 2006) (Fig 6).

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Bathonian–Callovian δ13CTOC datasets are scarce and usually of very low temporal

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resolution (Kenig et al., 1994; Nunn et al., 2009; Jach et al., 2014; Koevoets et al., 2016). Nunn et al. (2009) noted a decoupling event between δ13Ccarb and δ13CTOC during the Early Callovian in

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Scotland (within a lithological unit containing Kepplerites galilaei, Callomon in (Morton and

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Hudson, 1995)) and recoupling in the Middle Callovian positive CIE. Nunn et al. (2009) stated: “it may relate to either a reduced exchange of carbon between the oceans and the atmosphere (for reasons unknown) or to the presence of local noise in the δ13C record”. This dataset is likely younger than the studied interval of the Lusitanian Basin. In Portugal, the δ13Ccarb-TOC minimum recorded around the Julii/Histricoides subzones (Retrocostatum Zone) is interpreted as the δ13C minimum interval of the Early–Late Bathonian δ13Ccarb negative CIE (Figs 4 and 6). The δ13Ccarb-TOC decoupling event is dated from the Late Bathonian (Retrocostatum Zone, Histricoides Subzone)–Early Callovian? (Mangold, 1990). The

Journal Pre-proof termination of the decoupling event is materialised by the Early Callovian δ13CTOC positive CIE, associated with the later stages of the Late Bathonian–Early Callovian δ13Ccarb positive CIE. In Portugal, the Late Bathonian–Early Callovian δ13Ccarb positive CIE has an amplitude of about 1.5–2 ‰ (Fig. 6). Based on available ammonite data from Portugal (Mangold, 1990), the δ13CcarbTOC

decoupling event is presumed to have lasted ~ 0.3 Myr (Ogg et al., 2016).

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5.4. Possible mechanisms leading to the Late Bathonian–Early Callovian δ13Ccarb-TOC decoupling

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event in the Lusitanian Basin

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Covariant δ13Ccarb-TOC is usually interpreted as evidence that both carbonates and SOM have retained their original δ13C composition (Oehlert and Swart, 2014). On the other hand,

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decoupled δ13Ccarb-TOC records are much more infrequent and usually remain unexplained (Nunn

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et al., 2009; Yager et al., 2017) or are taken to reflect i) diagenetic alteration (Silva et al., 2013),

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ii) detrital flux of exogenous (fossil) OM or migration of hydrocarbons (Johnston et al., 2012), iii) local changes in DIC or DOC (Patterson and Walter, 1994; Immenhauser et al., 2003; Suan et

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al., 2015), or iv) the Rothman ocean model (changes in size of the OOC pool and non-steady

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state C cycle, see the Introduction section) Rothman et al., 2003) and variations thereof (Fike et al., 2006; Peltier et al., 2007; Swanson-Hysell et al., 2010). As discussed previously, the temporal evolution of the Upper Bathonian–Lower Callovian paired δ13Ccarb-TOC record from the Lusitanian Basin is here interpreted to reflect environmental, oceanographic, and diagenetic conditions and variation contemporaneous with deposition. The lack of correlation between δ13Ccarb and δ18Ocarb (Fig. 5) and the similarity of our δ13Ccarb record with several contemporaneous datasets suggest a weak diagenetic imprint in the δ13Ccarb record and a temporal, at least, regional variation pattern (Fig. 6). Extensive diagenetic

Journal Pre-proof alteration or terrestrial “contamination” are discarded as major mechanisms driving the δ13CTOC trend (see section 5.2). Explaining the observed invariant δ13CTOC record using the basinal OM vs exogenous OM model (Johnston et al., 2012) also seems unlikely. To our knowledge, there is no evidence of an exogenous source of organic carbon in the Lusitanian Basin, either from the erosion of organic-rich shales on land or hydrocarbons seepage. Provokingly, and as a working hypothesis, we suggest that the δ13Ccarb-TOC decoupling

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event and the Bathonian–Early Callovian CIEs resulted from the addition and removal of carbon

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from the OOC reservoir, likely DPOOC (Figs 1, 6, and 7). A similar mechanism can be evoked

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to explain the likely younger decoupling event and the Middle Callovian δ13Ccarb positive CIE in

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Nunn et al. (2009). As in Rothman et al. (2003), it is speculated that the mostly invariant δ13CTOC record during the decoupling event resulted from the accumulation of a large OOC reservoir,

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which increased the residence time of carbon in the oceans and buffered δ13CTOC (slow

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equilibrium) to changes in δ13CDIC (see Ridgwell and Arndt (2015) regarding the limitations of this interpretation). The relative size, chemical and isotopic composition, and oceanographical

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distribution of the DPOOC and DIC pools and composition and magnitude of the fluxes between

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the different superficial carbon reservoirs, including external inputs, are unknown (Figs 1 and 7) and limit further exploitation of proposed process in the Mesozoic. Investigation of contemporaneous sections from other geographical location, development of methods for a complete and detailed characterisation of SOM, and development of carbon cycle models that explicitly include OOC and DPOOC as a readily exchangeable superficial carbon reservoir are needed to constrain spatially and temporally the Bathonian–Callovian intervals. The accumulation of OOC implies a non-steady stated of the short-term carbon cycle in the Bathonian–Callovian.

Journal Pre-proof However, and regardless of the Bathonian–Callovian oceans ability to accumulate the vast amounts of OOC required to effectively buffer δ13CTOC (see Ridgwell and Arndt (2015) regarding practical limitations of this interpretation), the temporal variation in the size of a more “modern-sized” OOC reservoir may have had impacted δ13Ccarb and δ13CTOC, either via oxidation of DPOOC to CO2 or enhanced burial of organic matter from OOC (Rothman et al., 2003; Fike et al., 2006; Sexton et al., 2011; Ridgwell and Arndt, 2015). In the following paragraphs, it is

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assumed that the accumulation of OOC and DPOOC occurred in a large area of the global ocean.

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A testable prediction of our working hypothesis concerns the fate of the DPOOC pool

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and its impact on the global carbon cycle and climate: if part of this reservoir is oxidized to CO2, increases in atmospheric pCO2 and temperature are likely to occur and a negative CIE and

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recoupling of δ13Ccarb-TOC are expected; if DPOOC is exported to sediments, generalized organic

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matter deposition may occur and a positive CIE and recoupling of δ13Ccarb-TOC are expected (see

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also (Rothman et al., 2003; Fike et al., 2006; Bekker et al., 2008; Sexton et al., 2011; Ridgwell and Arndt, 2015)) (Fig. 7). The Early–Late Bathonian δ13Ccarb negative CIE (about 0.5 ‰ in the

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La Cornicabra section, about 1.5–2 ‰ in Portugal) and the Early Callovian δ13Ccarb negative CIE

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(about 0.9 ‰ in the La Cornicabra section) are here explained as resulting from the oxidation of DPOOC to CO2 (Fig. 6 and 7). The causes for the different amplitudes of the δ13Ccarb negative CIE between Portugal and Spain may be related to the poor resolution of the La Cornicabra section, diagenesis artefacts, spatial heterogeneity of the OIC and OOC reservoirs, different carbonate sources, or local factors affecting δ13Ccarb. Using the modern ocean carbon inventory as a reference and a very simplistic mass calculation exercise with an initial isotopic composition of δ13CDIC of 2 ‰ and an DPOOC source isotopic signature of -27 ‰ (eq 1.3b in Ridgwell and Arndt, 2015), negative CIEs with amplitudes of 0.5, 1, and 2 ‰ can be explained by the

Journal Pre-proof oxidation of about 700, 1500, and 3000 Pg of C from DPOOC respectively. These quantities represent roughly one to four times the modern-day DOC inventory (Hansell, 2013). As predicted, and despite the lack of detailed records, the two negative CIEs are apparently associated with relatively higher paleotemperatures (Dera et al., 2011). The Early Callovian δ13CTOC positive CIE and recoupling with δ13Ccarb (coeval with δ13Ccarb maximum of the latest Bathonian–Early Callovian δ13Ccarb positive CIE, O’Dogherty et

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al., 2006) is interpreted to be the result of increased efficiency in OOC and DPOOC export into

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sediments, materialized by the generalized deposition of organic matter during the Early

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Callovian (Fig. 6 and 7). Lower Callovian marine and terrestrial organic-rich facies and black

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shales are known from Spain (Bullatus Zone, Lower Callovian (Gräfe, 2005)), Northeast Scotland (Macrocephalus zone, Lower Callovian (Nagy et al., 2001)), central North America

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(Lower Callovian (Imlay, 1981)), Himalayas (Discus?–Gracilis zones; Upper Bathonian–Lower

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Callovian (Yin, 2007)), east Greenland (Lower Callovian (Callomon, 1993)), Atlantic Canada (Lower Callovian (Mukhopadhyay and Wade, 1990)), west-Gulf of Mexico (Cantú–Chapa,

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1969) and Japan (Kusuhashi et al., 2002; Handa et al., 2014).

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The wax and wane of the DPOOC component of the OOC reservoir are here viewed as the redistribution of carbon among Earth’s superficial reservoirs, driven by earth-orbital forcing mechanisms and internal ocean-atmosphere-biosphere-cryosphere feedbacks (Sundquist, 1985, 1991) and interaction with geosphere. Currently, there is a large uncertainty regarding sources and processes of accumulation and removal of recalcitrant DOC, which comprises the majority of the OOC reservoir in today’s oceans (Hansell, 2013; Follett et al., 2014; Legendre et al., 2015; Polimene et al., 2017; Wilson and Arndt, 2017). Examples of important sources of DOC include terrestrial contributions, viral lysis, phytoplankton release, “sloppy feeding” by metazoan

Journal Pre-proof grazers, egesta of protists and metazoan, particulate organic matter solubilization by bacterial and archeal ectohydrolases (Jiao et al., 2010; Hansell, 2013; Carlson and Hansell, 2015; Legendre et al., 2015). Oceanic accumulation of recalcitrant DOC is thought to occur by abiotic processes (via condensation reactions catalysed by light and metal complexation, polymerization, or adsorption to colloids) and biologically, mainly associated with the microbial carbon pump, limitation of heterotrophic bacterioplankton growth (via nutrient limitation, microbial

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community, pressure, and temperature), and eukaryotic production; the main processes of

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removal include biotic consumption by heterotrophic prokaryotes (via bacterioplankton

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consumption or remineralization) and eukaryotes, photochemical oxidation of recalcitrant DOC

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via ocean mixing, sorption of DOM onto sinking particles, condensation of gels, and hydrothermal circulation (see, for example, (Druffel et al., 1992; Jiao et al., 2010, 2014; Carlson

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et al., 2011; Arndt et al., 2013; Hansell, 2013; Follett et al., 2014; Legendre et al., 2015;

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Ridgwell and Arndt, 2015; Carlson and Hansell, 2015; Dittmar, 2015; Cael et al., 2017; Wilson and Arndt, 2017) and references therein).

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Any of the aforementioned processes, or a combination thereof, might have modulated

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the changes in size of the OOC reservoir and the DPOOC component in the Bathonian– Callovian. For example, it was suggested that ocean temperature changes might affect the efficiency of the biological carbon pump and the partitioning of organic carbon between POC and DOC, with warming shifting the partitioning between POC and DOC toward enhanced accumulation of dissolved compounds (Wohlers et al., 2009; Legendre et al., 2015). It seems reasonable to extrapolate that astronomically paced variation of several paleoclimatic parameters may have had resulted in changes of the relative efficiency of the biological carbon pump and microbial carbon pump and in the postulated changes in the sizes of the several superficial

Journal Pre-proof carbon reservoirs (Fig. 7). In addition to the effect of temperature in the biological carbon pump, astronomically paced variation in insolation might have also controlled the overall rates of photochemical oxidation of recalcitrant DOC and nutrient availability via climate and weathering. Even if we are still distant from a reliable model capable of describing the variations in size of the different oceanic and atmospheric carbon reservoirs in short-lived geological

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intervals, the working hypothesis of accumulation and remineralization of OOC and DPOOC as

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a first-order control on the carbon cycle and climate during specific intervals of the Mesozoic

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provides a testable hypothesis to explain the astronomically paced short-lived ~1 ‰ negative

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CIEs that infuse the Mesozoic, especially the small fluctuations within larger-scale CIEs associated with oceanic anoxic events (Schlanger and Jenkyns, 1976; Hesselbo et al., 2007).

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Small-scale isotopic perturbations that might be explained by the OOC working hypothesis can

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be found, for example, in the Early Jurassic Liasidium Event (Hesselbo et al., 2019), Early Toarcian oceanic anoxic event (OAE) (Hesselbo et al., 2007; Boulila et al., 2014; Thibault et al.,

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2018), or Late Aptian–Early Albian OAE 1b cluster (Herrle et al., 2004; Föllmi et al., 2006). If

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OOC, either alone or combined with additional inputs from the terrestrial biosphere, marine hydrates, and/or enhanced volcanic outgassing, can explain larger amplitude Mesozoic CIEs must also be qualitatively explored (Ridgwell and Arndt, 2015).

6. Conclusions and final remarks In this study, we present evidence of at least one δ13Ccarb-TOC decoupling event in the Bathonian–Callovian. Provokingly, we suggest that this event resulted from the accumulation of a large pool of oceanic organic carbon (OOC), likely dominated by dissolved and fine particulate

Journal Pre-proof oceanic organic carbon. Expansion of the OOC reservoir increased the residence time of carbon in the oceans and buffered δ13COOC (approximated by δ13CTOC) to changes in δ13C of oceanic inorganic carbon, dominated by dissolved inorganic carbon (approximated by δ13Ccarb). Variation in the size of the OOC pool and redistribution of carbon among the superficial carbon reservoirs provides an alternative explanation for the carbon isotopic excursions (CIEs) observed in the Bathonian–Callovian; oxidation of OOC to CO2 resulted in small-scale negative CIEs (~ 1‰),

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whereas the change from OOC production and accumulation to carbon export to sediments lead

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to positive CIEs. The reduction in size of the OOC pool lead to the recoupling of the δ13Ccarb-TOC

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records. Other small-scale (~ 1‰) δ13Ccarb-TOC excursions in the Mesozoic may have resulted

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from similar episodes of OOC accumulation and oxidation.

We hope that our bold claims stimulate further research in this often-overlooked interval

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and highlight the need to explicitly consider OOC (especially DOC) in current carbon cycle

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modelling efforts of high-resolution geological records. The working conceptual model presented here for the Bathonian-Callovian is readily testable by the acquisition of other detailed

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paired δ13Ccarb-TOC records and comprehensive characterization of organic matter source from

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other basins and detailed modelling of the carbon cycle with an explicit consideration of the short-term cycle and its major carbon reservoirs.

Acknowledgements This research was supported by the Source Rock and Geochemistry of the Central Atlantic Margins consortium (Dalhousie University, Basin and Reservoir Lab, PI – G. Wach). L. V. Duarte was supported by FCT, through the strategic project UID/MAR/04292/2019 granted to the Marine and Environmental Sciences Centre (MARE, Portugal). R. L. Silva was also partially

Journal Pre-proof supported by the Irish Centre for Research in Applied Geosciences, iCRAG (project: Temporal and spatial variability in Lower Jurassic hydrocarbon source rock quality in Irish off-shore marine basins, PI - Micha Ruhl). We also thank Katja Fennel, Stephen P. Hesselbo and Micha Ruhl for helpful scientific discussions on some of the topics covered in this manuscript; and the Editor-in-Chief Fabienne Marret-Davies and two anonymous referees for their insightful

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comments that greatly benefited the manuscript.

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Journal Pre-proof Figure captions Figure 1. a) Cartoon illustrating the main oceanic carbon reservoirs and selected transfer processes (Emerson and Hedges, 2008; Jiao et al., 2010, 2014; Hansell and Carlson, 2013; Carlson and Hansell, 2015; Ridgwell and Arndt, 2015); b) Schematic depiction of the sizes of aggregates containing organic carbon in the ocean (modified from Verdugo et al., 2004; Emerson and Hedges, 2008). 1 - air-water gas exchange; 2 - respiration; 3 - primary production

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(photosynthesis); 4 - resuspension/erosion/influx/oxidation; 5 -microbial biomass (heterotrophy);

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6 - phytoplankton production, exudation, grazing, viral lysis, solubilization from POC, riverine

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input; 7 - respiration; 8 - respiration/oxidation; 9 - dissociation; 10 - carbonate production and precipitation; 11 - carbonate dissolution; 12 - dissolution; 13 - sedimentation; 14 – microbial

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activity; 15 – phototransformation; 16 - sedimentation (PIC)/precipitation (DIC); 17 –

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methanogenesis.

Figure 2. Simplified geological map of the central-northern part of the Lusitanian Basin

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(Portugal) and distribution of the carbonate units of Lower Jurassic age.

Figure 3. Stratigraphic units and main sedimentological features of the Lower and Middle Jurassic series in the Lusitanian Basin (Portugal) (modified from (Azerêdo et al., 2020). Major Transgressive-Regressive and Transgressive-Regressive facies cycles from (Azerêdo et al., 2014). Sedimentary discontinuities from Soares and Duarte (1997).

Figure 4. Stratigraphic log, TOC, and kerogen and carbonate δ13C record of the Late Bathonian– Lower Callovian boundary in the Cabo Mondego Formation, Lusitanian Basin (Portugal).

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Figure 5. δ13C versus δ18O cross plot of the analysed samples from the Late Bathonian–Lower Callovian boundary in the Cabo Mondego Formation, Lusitanian Basin (Portugal).

Figure 6. Stratigraphic log, TOC and carbonate δ13C record of the Middle Bathonian–Lower Callovian boundary in the Cabo Mondego Formation, Lusitanian Basin (Portugal) and

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comparison with the δ13C record of several other European basins; Umbria-Marche Basin

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(Bartolini et al., 1996); Southern Alpes (Jenkyns, 1996), External Subbetic Basin (O’Dogherty et

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al., 2006) (LOESS smoothing, smoothing factor = 0.2, confidence interval = 95%).

Figure 7. Cartoon illustrating the relative sizes and variation of the several carbon pools in the

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oceans in the Bathonian–Callovian. The relative size, chemical and isotopic composition, and

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oceanographical distribution of the DPOOC, DIC, and atmospheric CO2 pools and composition and magnitude of the fluxes between the different superficial carbon reservoirs, including

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external inputs, are unknown.

Journal Pre-proof Highlights

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At least one δ13Ccarb-TOC decoupling event occurred in the Late Bathonian–Callovian Decoupling resulted from the accumulation of oceanic organic carbon (OOC) The enlargement of the OOC pool increased C residence time and buffered δ13CTOC to change The wane and wax of the OOC pool resulted in carbon isotopic excursions These decoupling events are evidence a non-steady state carbon cycle

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